Keywords: volcanic eruptions, orthopyroxenes, zonations, timescales, magma dynamics Scanning Electron Microscope ....

« Réalise-le ce rêve ! » Inscription vue dans les rues de Hyères, été 2022

REMERCIEMENTS

Dear rapporteurs and members of my PhD jury, I kindly thank you again for accepting to take part in my jury and for your time to evaluate my work: Professors Luca Caricchi, Chiara Maria Petrone, Fidel Costa, Caroline Martel and Daniele Morgavi. Ce travail de thèse a été possible grâce aux financements par l'ANR V-Care, par l'Institut de physique du globe de Paris et l'Université Paris Cité. Merci de m'avoir donné ce cadre de travail idéal pendant ces 3 ans et plus.

Voici donc le moment tant attendu, la rédaction des remerciements ! Comme c'est la première chose que je lis dans la thèse de quelqu'un d'autre, je vais essayer d'être à la hauteur et de trouver les mots justes.

Nous voici déjà en octobre 2022, 3 ans et un mois après notre retour de mission au Kamchatka, la terre de feu et de glace, par laquelle ma thèse volcanique a commencé ! Ainsi, je tiens à remercier tout d'abord mes directeurs de thèse, Georges et Hélène, grâce à qui tout cela a été possible. Merci de me suivre depuis maintenant 4 ans, depuis les cours de Systèmes volcaniques, puis durant le stage de Master 2 et le stage en Guadeloupe (on pourrait même remonter en décembre 2016 où, Georges, tu étais venu faire une conférence pour présenter ton métier de volcanologue à l'ENSG !). Merci à vous deux de m'avoir fait confiance et donné l'opportunité de faire cette thèse, sur la Dominique et le Kamchatka (région où même les Kraftt ne sont pas allés, m'a dit Audrey un jour). Je me souviendrai toujours de ce premier tour d'hélicoptère, l'émerveillement aux pieds du Bezymianny et du Kizimen, ces marches pour récolter nos précieux échantillons quand la météo nous le permettait et les thés sous la tente pendant que froid et brouillard régnaient dehors, la journée riche en émotions avec les Brothers et l'ours que l'on n'a pas vu avec Hélène et Caroline… Of course! Je voulais aussi vous remercier pour tout ce que vous m'avez appris et pour tout le temps que vous m'avez consacré, le temps passé à relire mes fichiers, discuter des données et des figures (les PPR !!). Merci d'avoir toujours été disponibles, même pendant le week-end, merci pour votre efficacité, vos doubles corrections. Merci de votre pertinence et exigence, patience et passion de la volcanologie, qui m'ont motivée (et me motivent toujours) à essayer de faire au mieux. J'ai vraiment énormément appris à vos côtés (même si c'est toi Georges qui cassait encore au début de ma 2 e année mes cailloux avant de les concasser !) et j'admire vos carrières. Je voulais aussi vous remercier pour tous les autres moments qui n'étaient pas forcément qu'autour du travail, pour tous ces repas partagés et discussions sur tout plein de sujets ; la mission nous a permis de tisser des liens forts, je vous en suis reconnaissante. Mille mercis.

Merci aussi à Edouard Kaminski et Etienne Médard d'avoir accepté de faire partie de mon comité de thèse, pour vos précieux conseils et discussions au cours de ces comités qui ont permis d'améliorer mes différentes études.

Many thanks to my co-authors and particularly to you, Dan, for working with me since my master's thesis and during my PhD, I learned a lot during our discussions.

Ensuite, j'aimerais remercier tous les membres de l'équipe des Systèmes Volcaniques à l'IPGP. Un grand merci en particulier aux habitués de la cantine, à Samia, pour ton aide surtout au début pour préparer les échantillons, pour ta bonne humeur et les séances de sport plus ou moins assidues avec Nadia et Marie ; Pierre pour ta bienveillance ; Arnaud, voisin du Sud, avec qui j'ai pu débuter dans le monde de l'enseignement au début de ma thèse ; Pascale avec qui on a passé un super stage terrain en Auvergne et Telly pour ta patience et pour toutes les questions administratives. Merci aussi à Marina, Jean-Christophe, Fidel, Anne et François pour vos encouragements. Un grand merci aussi à Aline de m'avoir fait visiter l'observatoire du Piton de la Fournaise quand nous étions de passage sur l'île intense avec Aurélien, c'était vraiment génial.

Et puis, bien sûr, MERCI aux Volcano Youngsters et co-bureaux du 469 ! First, Abigail, thank you for all the happy moments, shared meals either in Montrouge or in Paris and for your precious help! Giulia pour ta douceur et toutes nos discussions (forza pour la fin !!), Mathilde pour les petits moments partagés pendant la thèse. Puis merci à ceux arrivés après, à Raphaël (amitiés et pâté basque !), Chagnon, Laëtitia, la star de la sismo ! Merci pour tous tes petits messages et ton soutien ; j'adore toutes nos pauses reboostage/gossip/émerveillement de belles photos de volcans ou de girafes. Et merci aussi à Audrey, j'étais contente de t'avoir recroisée pendant la fin de thèse ! Désolée de vous avoir tourné littéralement le dos pendant ces 3 ans mais j'ai adoré partager ce bureau avec vous, ces repas, pauses café/bars et moments de rigolades (au final, j'étais le plus souvent tournée vers vous !!). Merci à toi aussi Olivier, tu es arrivé au cours de ma dernière année mais j'ai l'impression de te connaître depuis longtemps (c'est le réseau Géol ça haha), merci d'avoir répondu à mes questions sur la CSD et pour tous tes conseils ! Rotorua nous voilà ! Merci à Stephan aussi, pour ton aide au MEB, tes beaux calendriers et surtout pour tes blagues qui ont éclairé pas mal de journées (chef d'oeuvre).

Merci ensuite à mes amis du 4e ou des autres étages, pensées particulières à Clément et tes aventures rocambolesques (au Moz !!) and Lukas (dankeschön for this initial crew with Clément, for the parties and fun we had during these 3 years). Merci à Alexandra, ma binôme de chargées de TPs cailloux ! On en a fait du chemin depuis le master. Toutes les petites pauses partagées avec toi font toujours plaisir ! Merci aussi à Gaëlle (pour les déjeuners dans le jardin des plantes ou en Guadeloupe ; vive les racoons !), à Sophie pour ton soutien mais aussi à Alexandre, Francine, Charlotte et tous les autres.

Ensuite, merci aux membres de l'ISTeP, ma deuxième maison au MEB ou à la sonde !! Choukran bezef à toi Omar pour ton aide au MEB et pour les discussions nostalgiques sur le Maroc. Michel et Nicolas, mille mercis pour votre aide durant ces 3 ans et + (ça a bien bouclé !!) et pour toutes les discussions en attendant les calibrations. Ma famille en aura entendu parler de ces séances de 10 000 points où je devais absolument finir de programmer avant de rentrer ! Merci aussi à Eric pour toutes les lames minces et les échantillons de Perugia, je ne sais pas comment j'aurais fait sinon pour qu'ils soient coupés correctement ! Merci aussi à toi Thiebaut, binôme de terrain, pour ta bonne humeur et d'avoir toujours été disponible pour discuter de tout et de rien. Merci aussi à Adeline pour ton énergie (vive Dua !!) et Maylis pour ton soutien et ta gentillesse.

Thank you to the Kamchatka team on the field for making it a unique and amazing experience, Caroline, Saskia, Pascal pour tous les bons moments sur le terrain et les discussions sous la kitchen tent; thank you also to Anna, Sasha, Marina, Oleg, Ilya, Andreas and René.

Grazie anche al Dipartimento di Geologia di Perugia per avermi ospitato per queste 5 settimane. Daniele and Diego, thank you for accepting to start this collaboration, I learned so much this year with you, as well as with Fabio (by Skype!). Thank you also to Maurizio, Max, Alessandro, Lisa and Monica. I hope that this collaboration is the first step of many to come. It was a pleasure to meet you all, to melt some rocks and to live in beautiful Perugia for these few weeks, to enjoy the Italian way of life for a bit! J'en viens à remercier mes amis du LVC, Mallorie, Axel, Théo, Elliot, Mathieu, Alexandre et surtout Gabrielle pour tes petits messages pendant les derniers mois très intenses et pour tous ces moments passés ensemble. Merci aussi à mes amis de prépa, particulièrement à Thomas, Jerem et Olívia ! À mes copains de Nancy, Elise, que cela soit à Paris, en Mayenne ou dans le Sud, ça me fait toujours trop plaisir de te voir ; Paul, William, la team de stage terrain du love, toujours partants pour des lasagnes haricots blancs et lardons ; Nathanaël, Anaïs à La Réunion et tous les autres (Sido et Tatiana !). Merci aussi à mes amis de master de l'IPGP (dont Rachid, Cyril et Nathan expatrié en Bolivie) ! Merci aussi Félicien, pour les aventures à La Réunion avec Aurélien il y a 2 ans, c'était trop cool (Coco à fond dans la voiture !!). Je te souhaite plein de bonnes choses pour ta thèse et la suite !! Pour tous ceux qui ne sont pas explicitement cités, je pense à vous, un grand merci à tous pour vos messages et pensées pendant la thèse, et à ceux qui sont partis mais à qui je penserai toujours, Razane qui est partie brutalement au cours de ma première année et mes deux arrières grand mères.

Enfin, tout cela n'aurait pas été possible sans le soutien continuel de ma famille et de mes proches. Merci à mes grands-parents pour vos appels et messages, j'adore toujours autant venir vous voir au Grand Clou et à Giens. Merci à Cécile, Vincent, Gaëlle et Arnaud et mes cousines pour toutes vos attentions et pensées. Merci aussi à Sabine et Gioacchino pour leur soutien, ainsi qu'à Baptiste et Fanny. Et une pensée pour ma petite Mimi… Je te retrouve un peu à travers notre Maïdo, même s'il est encore plus foufou ! Et surtout, merci infiniment à mes parents d'avoir toujours cru en moi et d'avoir toujours été là pour moi. Merci pour cette vie incroyable que vous nous avez offerte avec Elaura à l'étranger et après, qui nous a façonné et donné le goût du voyage. Vous êtes mes plus grands modèles. Merci pour tous les petits moments partagés à votre retour d'Italie cet été qui m'ont donné plein d'énergie pour le dernier mois, pour les cours de dessin partagés avec toi, Maman, les petits plats thaïlandais de toi, Papa, et pour tous vos appels, attentions, messages et votre aide pour la relecture finale !! Merci maintenant à ma sis', pour tout et tous les fous rires (sympa la robe !) et de toujours me donner du courage pour "mon taff". Trop hâte de partir en road trip ! Et pour finir, Aurélien, tu sais déjà tout mais merci de m'avoir aidée à tenir jusqu'au bout, d'avoir toujours été là. Merci pour ce rêve éveillé que l'on vit depuis maintenant presque 8 ans.

Léa, le 2 octobre 2022

Dynamique spatio-temporelle des systèmes de stockage magmatique, vers une "horloge pré-éruptive". Application à l'île de la Dominique (Petites Antilles) et au Kamchatka (Russie)

RÉSUMÉ

Les éruptions volcaniques font partie des événements les plus dévastateurs sur Terre. En contexte de subduction, les éruptions mettent généralement en jeu des magmas différenciés : après de longs séjours dans des réservoirs magmatiques, les magmas basiques issus de la fusion partielle du manteau évoluent vers des compositions différenciées, donc plus riches en silice, plus visqueuses et plus riches en éléments volatils. Ces magmas peuvent ainsi donner des styles éruptifs variés plus ou moins explosifs qui dépendent de différents paramètres dont la composition du magma, le volume de magma mis en jeu, et le comportement des éléments volatils lors de la remontée dans les conduits d'alimentation.

Réussir à estimer le temps séparant les changements de conditions de stockage des magmas à l'intérieur d'un réservoir magmatique et le début d'une éruption est l'un des défis majeurs en volcanologie. Une observation directe du système de stockage des magmas sous les volcans est impossible mais cette échelle de temps peut être accessible, notamment grâce à l'étude de cristaux amenés en surface par une éruption volcanique passée. En effet, comme les cernes des arbres, qui peuvent permettre d'obtenir des informations sur des événements passés, les cristaux qui grandissent au sein de réservoirs magmatiques alimentant des volcans enregistrent les modifications des conditions de stockage des magmas avant une éruption. Pour beaucoup d'éruptions, il a été montré que des réinjections de magma ou des événements de décompression pouvaient se produire dans le réservoir et être enregistrés par les cristaux sous la forme de zonations (i.e. zones internes ou externes de compositions différentes du coeur du cristal). Avec le temps, un rééquilibrage des compositions chimiques entre les zonations par la diffusion d'éléments se produit (diffusion du fer et magnésium pour les orthopyroxènes, diffusion du titane dans les magnétites, par exemple). Lors d'une éruption, la diffusion est figée par le refroidissement rapide des magmas émis et il est alors possible de déterminer le temps écoulé entre les processus magmatiques ayant formé les zonations dans le réservoir et l'éruption.

Les temps de réactivation magmatique ont été étudiés au cours de cette thèse pour des volcans de zones de subduction, générant un volcanisme d'arc particulièrement dangereux, en se focalisant sur cinq éruptions explosives d'un volcan de l'île de la Dominique ainsi que sur deux éruptions récentes de la zone de subduction du Kamchatka (Russie). Des modèles spatio-temporels des dynamiques des zones d'alimentation magmatique à l'origine de sept éruptions sont donc proposés. Ce nouvel ensemble de données de temps de diffusion sur des volcans de zones de subduction permet une meilleure compréhension de ces systèmes, sachant que les études de temps de diffusion sur les volcans des Petites Antilles et du Kamchatka sont peu nombreuses, malgré des risques volcaniques importants pour le cas des Petites Antilles.

Pour l'un de ces derniers exemples, pour l'éruption de 2010 du volcan Kizimen au Kamchatka, les temps ont été corrélés avec les premiers signaux précurseurs de l'éruption (géophysiques et géochimiques) et permettront de mieux gérer les futures crises volcaniques sur ce volcan et plus généralement au niveau des zones de subduction. Ces résultats sur des éruptions passées de la Dominique et du Kamchatka avec la compréhension de la dynamique des réservoirs et les corrélations entre temps pétrologiques et signaux sismiques pour le Kizimen sont donc des avancées importantes. Si la zone de stockage des magmas pour un volcan donné est stable dans le temps, et si la corrélation entre les réajustements dans le réservoir et les signes précurseurs est montrée pour des éruptions passées, ces échelles de temps permettront, lors d'une réactivation de ces volcans, d'avoir une horloge prééruptive utile pour la protection des populations.

Mots-clés : éruptions volcaniques, orthopyroxènes, zonations, échelles de temps, dynamiques des magmas ABSTRACT Volcanic eruptions are among the most devastating events on Earth. In a subduction context, eruptions generally involve differentiated magmas: after long repose times in magma reservoirs, basic magmas resulting from the partial melting of the mantle evolve towards differentiated compositions, thus richer in silica, more viscous and richer in volatile elements. These magmas can give various eruptive styles more or less explosive depending on different parameters such as the composition of the magma, the volume of magma involved, and the behavior of volatile elements during the ascent in the volcanic conduits.

Estimating the time between changes in magma storage conditions within a magma reservoir and the onset of an eruption is one of the major challenges in volcanology. A direct observation of the magma storage system under the volcanoes is impossible but this timescale can be accessible thanks to the study of crystals brought to the surface by a past volcanic eruption. Indeed, like tree rings, which can provide information on past events, crystals growing within magma reservoirs feeding volcanoes record changes in magma storage conditions before an eruption. In many eruptions, it has been shown that magma injections or decompression events can occur in the reservoir and be recorded by the crystals as zonations (i.e. internal or external rims of different compositions than the crystal cores). With time, a homogenization of chemical compositions between zonations by diffusion of elements occurs (for example, the diffusion of iron and magnesium for orthopyroxenes or diffusion of titanium in magnetites). During an eruption, the diffusion is frozen by the rapid cooling of the emitted magmas and it is then possible to determine the time elapsed between the magmatic processes that formed the zonations in the reservoir and the eruption.

Pre-eruptive timescales have been studied during this thesis for subduction zone volcanoes generating particularly dangerous arc volcanism, focusing on five explosive eruptions from a volcano on the island of Dominica as well as two recent eruptions from the Kamchatka subduction zone (Russia). Spatiotemporal models of the dynamics of the magmatic feeding zones at the origin of these seven eruptions are thus proposed. This new set of diffusion timescales data on subduction zone volcanoes allows a better understanding of these systems, knowing that diffusion timescales studies on volcanoes of the Lesser Antilles and Kamchatka are few, despite the important volcanic risks for the case of the Lesser Antilles.

For one of these eruptions, for the 2010 Kizimen eruption in Kamchatka, the timescales were correlated with the first precursory signals of the eruption (geophysical and geochemical) which can allow to better manage future volcanic crises on this volcano and more generally in subduction zones. These results on past eruptions of Dominica and Kamchatka with the understanding of reservoir dynamics and correlations between petrological timescales and seismic signals for Kizimen are thus important advances. If the magma storage area for a given volcano is stable in time, and if the correlation between reservoir readjustments and precursory signals is shown for past eruptions, these timescales will allow, during a reactivation of these volcanoes, to have a useful pre-eruptive warning clock for the protection of the populations. [START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF]. …….45 Figure I.3: Melt inclusion entrapment in unzoned or zoned crystals and formation of the matrix or residual glass (adapted from [START_REF] Wallace | Volatiles in subduction zone magmas: concentrations and fluxes based on melt inclusion and volcanic gas data[END_REF]). ………………………………………………………...47 Figure I.4: Closure temperature for several elements in minerals and the potential of diffusion chronometry for a wide range of minerals (modified from Costa et al. (2021Costa et al. ( , 2020)), respectively). …49 2015)). …………………………………………...……………………………………………….56 Figure I.8: Diagram showing the time history of volcanic earthquakes recorded prior to and after an eruption (a) and run-up time (period between the unrest and the eruption) for explosive eruptions with a VEI of 4 or greater, since 1992 (b) (modified from Acocella (2021) and [START_REF] Cassidy | Explosive Eruptions With Little Warning: Experimental Petrology and Volcano Monitoring Observations From the 2014 Eruption of Kelud, Indonesia[END_REF], respectively). ………………………………………………………………………………………….58 Figure I.9: Example of a correlation between petrological and geophysical signals prior to an eruption (a) and relationship between run-up (time between the first signal of unrest and the eruption) and repose (time since the previous eruption) (b) (modified from Ruth et al. (2018) and Rassmussen et al. (2018), respectively). ………………………………………………………………………………………….60 Figure II.7: Chronology of the pumiceous eruptions of Dominica during the last 70 000 yr (modified from d [START_REF] 'augustin | Evidence for an Active, Transcrustal Magma System in the Last 60 ka and Eruptive Degassing Budget (H 2 O, CO 2[END_REF] and Boudon et al. (2017)). …………………………………………….80 Figure II.8: Location of Kamchatka. ………………………………………………….………………82

TABLE OF CONTENTS

LIST OF FIGURES AND SUPPLEMENTARY FIGURES

Figure II.9: Simplified map of Kamchatka and its volcanoes (Kamchatka Volcanic Eruption Response Team (KVERT), 2002) (a) and 3D visualizations (b-c). ….. ………………………………………….83 Figure II.10: Post-Eocene Kamchatka subduction zone history and geodynamic reconstructions (a1-a4) from 38 to 2 Ma (modified from [START_REF] Lander | The origin of the modern Kamchatka Subduction Zone[END_REF]). ………………………………………84

Figure II.11: Major structural elements and volcanoes in the CKD (modified from [START_REF] Koulakov | Mantle and Crustal Sources of Magmatic Activity of Klyuchevskoy and Surrounding Volcanoes in Kamchatka Inferred From Earthquake Tomography[END_REF]). …………………………………………………………………………………………...…...86 Figure II.12: Classification diagram for the Quaternary CKD lavas (Total Alkali Silica diagram, a) and (b) models for the magma sources (modified from Portnyagin et al. (2007a)). ……………………….88 Figure II.13: Representations of the magma reservoirs below the Klyuchevskoy group. …………….90 Figure II.14: Bezymianny volcano (modified from Shevchenko et al. (2020)). ……………………….91 Figure II.15: Bezymianny volcano, 1956 climactic phase of the eruption (modified from [START_REF] Shevchenko | The rebirth and evolution of Bezymianny volcano, Kamchatka after the 1956 sector collapse[END_REF]) and location of the deposits (modified from [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF]). ……………………………92

Figure II.16: Bezymianny volcano general eruptive timeline from 1956 to present-day. …………….94 Figure II.17: Cartoon model of physical processes prior to 1956, with three mixing components of Bezymianny evolving apart from each other via crystal fractionation (modified from [START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]). …………………………………………………………………………………. …………….97 Figure II.18: Seismic stations and Klyuchevskoy group. ……………………………………………. ..98 Figure II.19: Map of the deposits of the 2010-2013 eruption of Kizimen (image: NASA Operational Land Imager acquired on 2 September 2013) (modified from Auer et al. (2018)). …………………. ..100 Figure II.20: Pictures of the 2010 eruption of Kizimen and map of the deposits. ……………………101

Figure II.21: Network of the closest seismic stations of Kizimen volcano (modified from Sharikova and Firstov ( 2021)) (a-b) and seismic activity of Kizimen volcano from 01.01.1962 to 01.05.2011 (modified from Senyukov et al. ( 2011)) (c-f). ………………………………………………………………. ….102 Figure II.22: Kizimen volcano and timeline its pre-eruptive processes prior to the 2010 eruption with a model of the plumbing system (cartoons below the timeline are modified from Auer et al. (2018)). . ..103 Figure III.1: Pictures of the outcrops of Grand Fond (GF), Goodwill (GE) and PPR1 to 3 in Dominica (modified from Boudon et al. (2017) and d [START_REF] 'augustin | Evidence for an Active, Transcrustal Magma System in the Last 60 ka and Eruptive Degassing Budget (H 2 O, CO 2[END_REF]). …………………………………110 Figure III.2: Sketch maps of the eruptive deposits and samples localization. ………………………111

Figure III.3: Laterally directed blast deposits of the 1956 Bezymianny eruption (two layers visible in the photo: layer A and B) (photo by G. Boudon). …………………………………………………….112 Figure III.4: Three types of samples studied at Bezymianny. ……………………………………….112 Figure III.5: Kizimen and nature of the samples. ……………………………………………………113

Figure III.6: Measurement of densities and vesicularities (photos by S. Hidalgo). …………………114

Figure III.7: Preparation of samples from the rock sample to the crystal mounts of different types. ...115 Figure III.8: Main signals emitted during the interaction of an electron beam with a sample in the SEM.

………………………………………………………………………………………………….….…116

Figure III.9: Examples of SEM backscattered electron images of a thin section, rock mount and individual crystal images of Kizimen andesitic sample from the 2010-2013 eruption. ………………117

Figure III.10: Schematic representation of the crystallization of a crystal (here, an orthopyroxene, opx) from a liquid defined by a set of variables (pressure: P, temperature: T, composition: X) at time t1. . ..120 Figure III.11: Examples of defects in crystals (modified from Costa and Morgan, (2010)). …………121

Figure III.12: Notations used in the mass balance of fluxes for a rectangular solid (adapted from Costa et al. (2008)). Volume of a rectangular parallelepiped with each side having increasingly smaller lengths, e.g. dx, dy, dz. J(x) and J(x+dx) are the flows parallel to the x direction at the x and x+dx positions respectively. …………………………………………………………………………….…122 Figure III.13: Initial zoning profiles used in the diffusion time modelling (dotted lines) and their evolution after a given diffusion time (solid lines) (modified from Costa et al. (2008)). …………….124 Figure III.14: Illustration of the use of boundary conditions in closed (a-b) or open (c-d) systems (modified from Costa et al. (2008)). ………………………………………………………………….124 Figure III.15: Two analytical solutions of the diffusion equation for initial and boundary conditions encountered during modelling in natural crystals (one dimensional and without any dependence of D on concentration or position) (modified from Costa and Morgan (2010)). ………………………….126 Figure III.16: Arrhenius relationship between the reverse of temperature and the logarithm of the diffusion coefficient (example with values of diffusion coefficients for Ca diffusion in pyroxenes or pyroxenoids [START_REF] Freer | Null result" diffusion experiments with diopside: implications for pyroxene equilibria[END_REF]). ……………………………………………………………….……127 Figure III.17: Orthopyroxene solid solution and structure. ……………………………………….…130 Figure III.18: Comparison of diffusion times in orthopyroxenes using different interdiffusion coefficients. ……………………………………………………………………………………….…134 Figure III.19: Description of the diffusion time modelling method. …………………………………137

Figure III.20: Partial perspective of spinel (modified from [START_REF] Aragon | Cation diffusion in titanomagnetites[END_REF]). Oxygen atoms in red circles, cations in their initial positions (small empty circles) and cations in their final positions (solid). The arrows show two types of elementary jumps: between tetrahedra or from a tetrahedron to an octahedron. ……………………………………………………………………………………….….139 Figure III.21: From magnetite to Ti diffusion time. ………………………………………………....140 Figure III.22: Plumbing system of Etna (modified from Kahl et al. (2011)). …………………………141

Figure III.23: Crucibles containing the magmatic melts. ……………………………………….…. ..143 Figure IV.1: Dominica eruptions. ……………………………………………………………..……..151 Figure IV.2: The different zonations identified in the opx cargo: images and proportions. …………152 Figure IV.3: Frequency histograms of the En content of unzoned (a1-a5) and zoned opx of the five studied eruptions (single-zoned (SZ) opx: b1-b5 and multiple-zoned (MZ) ones: c1-c5). …..………153 Figure IV.4: Frequency histograms of the En content of zoned opx of the five studied eruptions. …. ..154 Figure IV.5: Frequency histograms of Fe-Mg interdiffusion timescales modelled in opx at 850 ± 25°C for Grand Fond (a, b) and 890 ± 9°C for core-rim and inner-outer rim boundaries of single-zoned (SZ) and multiple-zoned (MZ) opx of the two opx populations as in Figure IV.3 and 4,, PPR1 (e-f), PPR2 (g-h) and PPR3 (i-j). ………………………………………………………………155

Figure IV.6: Crystal System Analysis approach of the zonations recorded by the two populations of opx of the five eruptions studied (Grand Fond (a), Goodwill (b), PPR1 (c), PPR2 (d) and PPR3 (e)). ..…156 Figure IV.7: Simplified view of the storage areas and of the pre-eruptive dynamics of the magmas in the magmatic mush for Morne Trois Pitons-Micotrin. ………………………………………….……158 Figure VI.2.1: Kizimen volcano (Kamchatka). ………………………………………………………315

Figure VI.2.2: Examples of the timescales modelling in the zonations of plag of andesites (a-b) or in mgt (c-d). …………………………………………………………………………………………….322 Figure VI.2.3: SEM images showing typical zonations of studied crystals in dacite and andesite. ………………………………………………………………………………………………………..323 Figure VI.2.4: Histograms of compositions measured in opx and plag of the dacite (shades of blue) or andesite (shades of red). ……………………………………………………………………………..327 Figure VI.2.5: Mgt composition and plag compositions in minor elements of the dacite (shades of blue) or andesite (shades of red). …………………………………………………………………………..328 Figure VI.2.6: Histograms of the timescales histograms after modelling intracrystalline diffusion of (a) Fe-Mg in opx (dacite in blue and andesite in red) (all data from Ostorero et al. (accepted)), (b) Mg in plag and Ti in mgt with different colours corresponding to the plag types and zonations types in mgt (2 timescales on inner to outer rims in multiple-zoned mgt (MZ)). …………………………………….329 Figure VI.2.7: Uncertainties associated to the timescales in plag and mgt. ………. ..................…….330 Figure VI.2.8: Crystal System Analysis approach of the zonations recorded by the plag and opx for the dacite (a-b) and for the andesite (c-d). …………………………………………………………….…333 Figure VI.2.9: Cumulative number of the timescales estimated on the major compositional changes evidenced in plag, opx and mgt. …………………………………………………………………….336 Figure VI.2.10: Model of the reservoir dynamics during the recent history of Kizimen prior to the 2010 eruption recorded by plag, opx and mgt, on different timescales. …………………………………..337 Figure VI.3.1: Kamchatka peninsula and Kizimen volcano. ………………………………………..345 Figure VI.3.2: Andesite-dacite mixing textures of Kizimen 2010-2013 eruption samples shown at different magnifications, on a-b) macroscopic scale and c-d) backscatter electron (BSE) images. ...347 Figure VI.3.3: Schematic representation of the COMMA (not to scale, modified from Morgavi et al., 2015;[START_REF] Rossi | Exponential decay of concentration variance during magma mixing: Robustness of a volcanic chronometer and implications for the homogenization of chemical heterogeneities in magmatic systems[END_REF]. …………………………………………………………………………….349 Figure VI.3.4: Total Alkali Silica (TAS) diagram of all compositions: residual glasses in the crystal enclaves (cores, middle, margins and exterior areas), as well as felsic enclaves present in the crystallized enclaves, transects with or without crystals and end-members compositions. ………………………352

Figure VI.3.5: BSE images of plagioclase and oxides microlites in the experimental sample in the basaltic area (a) and natural samples (b-c, in the andesitic and dacitic domains, respectively) and microlite aspect ratios (d). ……………………………………………………………………………353

Figure VI.3.6: Crystal Size Distributions (CSD) of the experimental (basaltic area) and natural samples of Kizimen (in the andesitic and dacitic domains), derived by CSDCorrections 1.5 software [START_REF] Higgins | Closure in crystal size distributions (CSD), verification of CSD calculations, and the significance of CSD fans[END_REF][START_REF] Higgins | Measurement of crystal size distributions[END_REF]. ………………………………………………………………………………………….355 Figure VI.3.7: BSE images of the experimental samples. ……………………………………………356

Figure VI.3.8: Compositional variation along three representative transects in the glasses (a1-d3) for major elements: a) SiO2, b) CaO, c) MgO, d) FeO. The Supplementary Figure VI.3.8 shows the locations on BSE images of the transects a1-d1 in the experimental slice. Compositions of initial basaltic and dacitic melts are marked (blue and yellow, respectively), with error bars including analytical uncertainties. a4-d4) Compositional histograms of the compositions of the transects without crystals and with crystals. …………………………………………………………………………………….359 Figure VI.3.9: Harker diagrams for the glasses (in crystal enclaves, felsic enclaves, transects) and endmembers compositions. ……………………………………………………………………………. ..360 Figure VI.3.10: Experimental plagioclase crystals and their compositions. …………………………361 Figure VI.3.11: Schematic sketch of the processes deciphered here, the kinetics of the first mixing event related to the injection of a basalt in the dacite, leading to the formation of an andesite (a) and the timescales of the third magma mixing event that occurred very shortly prior to the eruption of Kizimen (few minutes to hours; c). ……………………………………………………………………………362 Supplementary Figure images of the locations of the transects done in Figure VI.3.8: a) through filaments with crystals; b) through a felsic enclave in a crystal enclave and c) through filaments without crystals near a crystal enclave. …………………………………………..375 Figure VI.4.1: Natural basaltic andesite and dacite (a-b) and of the same ones after the melting process (c-d). …………………………………………………………………………………………………379

Figure VI.4.2: BSE overview images of the two representative experimental slices, which shows a dark grey background (the felsic domain), small (< 10 µm) or bigger bubbles (20-30 µm) and wide light grey areas. …………………………………………………………………………………………………380

Figure VI.4.3: Filaments textures and presence of crystals. Stretching and folding of the filaments is observed when crystals are present. ……………………………………………………………….…381 Figure VI.4.4: Details on filaments and crystals in small crystal enclaves or in bigger filaments. Stretching and folding dynamics are visible in the filaments (a-b). ………………………………….382 Figure VI.4.5: Details of a crystallized enclave. Mafic filaments are visible on the top right-hand corner, through which the transect was performed. ………………………………………………………….383 Figure VI.4.6: Total Alkali Silica (TAS) diagram of all compositions: residual glasses in the crystal enclaves (cores, middle, margins and exterior areas), as well as transects with or without crystals and end-members compositions. …………………………………………………………………………383

Figure VI.4.7: Harker diagrams for the glasses (in crystal enclaves and transects) and end-members compositions. ………………………………………………………………………………………. ..384 Figure VI.4.8: Compositions of the experimental plagioclase crystals. a) Compositional histogram of the An# content of the cores, rims and edges of the plagioclase crystals; b) Ternary diagram of the plagioclase [START_REF] Marshall | Ternplot: An excel spreadsheet for ternary diagrams[END_REF]. ………………………………………………………………………..385 Figure VII.1: Compilation of pre-eruptive and syn-eruptive diffusion timescales in ascending order modelled in several types of minerals in different volcanic systems, involving mafic or evolved magma compositions. ………………………………………………………………………………………...396 Appendix Figure A.1: Dominica and studied ignimbritic eruptions. ………………………………...461

CHAPTER I: INTRODUCTION Résumé

Dans ce premier chapitre introductif, nous présentons un aperçu général sur les volcans et la genèse des magmas ainsi que sur les modifications des conditions de stockage que les magmas peuvent connaître pendant leur histoire pré-éruptive.

Les volcans sont généralement situés à la limite des plaques tectoniques, au niveau des dorsales océaniques, ou des zones de subduction mais également en domaine intraplaque : les volcans de points chauds, comme celui de Hawaï ou de La Réunion. Les volcans peuvent entrer en éruption, en émettant des gaz chauds, de la lave (magma qui a perdu ses gaz volcaniques) et/ou des téphras (cendres et roches volcaniques) éjectées. Au cours de l'histoire du volcan, des phases de construction/destruction de l'édifice volcanique peuvent se succéder. Environ 800 millions de personnes vivent à moins de 100 km d'un volcan et environ 1550 volcans terrestres ont été actifs au cours des 10 000 dernières années sur Terre (9 440 éruptions), l'observation de ces volcans est donc très importante (Pankhurst et al., 2018;United Nations, 2015). De plus, au cours du siècle prochain, les éruptions volcaniques à grande échelle sont cent fois plus probables de se produire que les impacts d'astéroïdes et de comètes [START_REF] Cassidy | Huge volcanic eruptions: time to prepare[END_REF][START_REF] Trilling | The Size Distribution of Near-Earth Objects Larger Than 10 m[END_REF] et il y a une chance sur 6 pour qu'une éruption de magnitude 7 ou plus se produise au cours de ce siècle [START_REF] Lin | Magnitude, frequency and climate forcing of global volcanism during the last glacial period as seen in Greenland and Antarctic ice cores (60-9 ka)[END_REF]. Or, le monde n'est pas préparé aux éruptions de fortes magnitudes, l'histoire des magmas et leurs dynamiques avant des éruptions volcaniques sont donc des processus importants à contraindre et à comprendre [START_REF] Cassidy | Huge volcanic eruptions: time to prepare[END_REF]. Comme le comportement volcanique passé et les processus pré-éruptifs des éruptions passées d'un volcan sont un guide précieux pour le comportement futur d'un volcan, l'étude de la stabilité ou non des zones de stockage des magmas, les cristaux et les liquides magmatiques émis par ces volcans sont étudiés (Pankhurst et al., 2018). Les processus pré-éruptifs sur des volcans spécifiques peuvent être comparés à des signaux géophysiques ou géochimiques détectés en temps réels sur le terrain avant les éruptions pour voir si ils opèrent sur une même échelle de temps (Pankhurst et al., 2018). L'observation en surface de ces signaux de réactivation, qu'ils soient géophysiques ou géochimiques, couplés aux données pétrologiques pourrait donner une indication du temps restant avant une éruption, la question du temps étant une problématique essentielle en volcanologie et en surveillance volcanique.

Les magmas proviennent de la fusion partielle de roches mantelliques ou crustales et sont constitués de cristaux, de liquides silicatés et de gaz volatils. Ils peuvent subir, durant leur histoire, différents phénomènes dont les principaux sont la cristallisation fractionnée, des mélanges de magmas ou une contamination par les roches environnantes de la croûte. Ainsi, la cristallisation fractionnée peut avoir lieu dans un réservoir magmatique, elle est caractérisée par la cristallisation de minéraux par la baisse de température du magma au cours du temps (Bowen, 1928). Durant cette évolution, le liquide évolue vers des compositions plus riches en silice et en éléments alcalins, ainsi qu'en éléments volatils (Bowen, 1928). Les magmas sont stockés dans la croûte sous les volcans. Au cours des dix dernières années, des modèles physiques suggèrent que les chambres magmatiques ne sont pas faciles à former et sont instables, contrairement à une vision ancienne de ces zones de stockage, caractérisée par la présence d'une chambre magmatique stable et dominée par la présence de liquide. Des travaux récents ont montré que les réservoirs pouvaient être présents sous la forme de plusieurs lentilles de magmas interconnectées ou non dans un mush magmatique (principalement constitué de cristaux mais aussi de liquide et d'éléments volatils, distribués de façon hétérogène dans l'espace et le temps) (Bachmann and Huber, 2016;Cashman et al., 2017;[START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF]. Cette vision pose de nouvelles questions à propos des zones de stockage des magmas et de leur évolution dans le temps.

Les hétérogénéités de composition dans les cristaux ou les zonations (bordures de compositions différentes du coeur du cristal) produites par différents phénomènes comme les mélanges de magma sont figés dans le temps après une éruption (Costa et al., 2020;Costa and Morgan, 2010). Ces réajustements correspondent à des changements de certains paramètres tels que la pression (P), la température (T), la fugacité en oxygène (fO2) (Costa et al., 2020(Costa et al., , 2008;;Costa and Morgan, 2010;Petrone and Mangler, 2021). Ces conditions de stockage peuvent être connues, d'une part, en analysant des traces des liquides volcaniques, piégés sous la forme d'inclusions dans les cristaux, s'étant transformé en verre lors de l'éruption par trempe chimique, et d'autre part, grâce à la pétrologie expérimentale. Cette dernière permet de reproduire au laboratoire l'assemblage des minéraux et les compositions des roches retrouvées en surface et d'en déduire les conditions de stockage des magmas. L'étude des verres matriciels ou résiduels, correspondants au liquide magmatique entourant les cristaux s'étant décomprimé et ayant plus ou moins microcristallisé lors de la remontée du magma et de l'éruption permet aussi d'accéder à la température du liquide magmatique et aux conditions d'ascension des magmas. Des changements des conditions de stockage magmatique se traduisent par la poursuite de la cristallisation des cristaux déjà présents dans le mush magmatique, dans des conditions différentes. Cette cristallisation entraîne la formation de bordures en périphérie des cristaux, qui ont une composition chimique plus ou moins différente de la composition initiale du cristal (zonations chimiques). Des cristaux zonés sont ainsi obtenus. Or, avec le temps, la diffusion intracristalline tend à rééquilibrer la composition entre les deux parties du cristal zoné. La vitesse de diffusion dépend de l'élément chimique, de la nature du cristal, de l'axe cristallographique considéré ainsi que de la T, P, fO2 (Ganguly and Tazzoli, 1994). La diffusion des éléments étudiés dans le minéral doit être assez rapide pour effacer partiellement la zonation mais assez lente pour que le cristal ne s'équilibre pas totalement aux nouvelles conditions (Costa et al., 2008). Ces cristaux zonés ou textures de déséquilibre peuvent être considérés comme des horloges cassées retrouvées sur une scène de crime, pour estimer les échelles de temps des processus antérieurs à une éruption [START_REF] Morgavi | Decrypting Magma Mixing in Igneous Systems[END_REF].

Les temps peuvent être modélisés dans différentes types de minéraux et contextes volcaniques, en prenant en compte différents éléments qui diffusent. Les temps modélisés dans plusieurs cristaux (olivines, plagioclases, pyroxènes, quartz…) montrent des échelles de temps allant de quelques minutes à des centaines d'années (Costa, 2021;Costa et al., 2020). Des âges radiogéniques peuvent aussi être mesurés et sont plutôt de l'ordre de la centaine à millions d'années. Une interprétation intéressante entre ces différences de temps est que la modélisation des temps de diffusion correspond aux dernier événements ayant perturbé le système (par exemple, la remobilisation du magma ou une intrusion magmatique) alors que les âges U-Th dans les zircons correspondent au temps que prend le magma pour s'accumuler (Costa, 2021;Costa et al., 2020;[START_REF] Turner | Measuring Timescales of Magmatic Evolution[END_REF]. Pour les éruptions récentes des volcans monitorés, il est intéressant d'étudier si ces temps pétrologiques obtenus par chronométrie de diffusion peuvent être corrélés entre eux avant plusieurs éruptions du même volcan, afin de déceler si des systématiques de temps de réactivation magmatique peuvent être identifiées. Ces études ont été réalisées pour plusieurs volcans mettant en jeu des magmas peu évolués (Albert et al., 2019;Kahl et al., 2011[START_REF] Kahl | Compositionally zoned crystals and real-time degassing data reveal changes in magma transfer dynamics during the 2006 summit eruptive episodes of Mt. Etna[END_REF]Longpre et al., 2014;[START_REF] Marti | Correlation of Magma Evolution and Geophysical Monitoring during the 2011-2012 El Hierro (Canary Islands) Submarine Eruption[END_REF]Pankhurst et al., 2018;[START_REF] Rae | Time scales of magma transport and mixing at Kīlauea Volcano, Hawai'i[END_REF]Sundermeyer et al., 2020a) ou pour des volcans d'arcs de subduction [START_REF] Cooper | [END_REF][START_REF] Rasmussen | When does eruption run-up begin? Multidisciplinary insight from the 1999 eruption of Shishaldin volcano[END_REF]Ruprecht and Plank, 2013;Ruth et al., 2018). Peu d'études se sont intéressées aux dynamiques pré-ou syn-éruptives des magmas avant des éruptions de magmas évolués (Davydova et al., 2018;[START_REF] Kilgour | Timescales of magmatic processes at Ruapehu volcano from diffusion chronometry and their comparison to monitoring data[END_REF]Saunders et al., 2012;Utami et al., 2021). Les temps obtenus dans ces études sont de l'ordre de quelques jours à plusieurs mois à quelques années avant l'éruption et correspondent le plus souvent à des mélanges de magmas avant l'éruption ou d'injection magmatique (Albert et al., 2019(Albert et al., , 2016;;Costa et al., 2013;Davydova et al., 2018;[START_REF] Kahl | Compositionally zoned crystals and real-time degassing data reveal changes in magma transfer dynamics during the 2006 summit eruptive episodes of Mt. Etna[END_REF]Kahl et al., , 2011;;Longpre et al., 2014;[START_REF] Marti | Correlation of Magma Evolution and Geophysical Monitoring during the 2011-2012 El Hierro (Canary Islands) Submarine Eruption[END_REF]Pankhurst et al., 2018;[START_REF] Rae | Time scales of magma transport and mixing at Kīlauea Volcano, Hawai'i[END_REF][START_REF] Rasmussen | When does eruption run-up begin? Multidisciplinary insight from the 1999 eruption of Shishaldin volcano[END_REF]Ruprecht and Plank, 2013;Ruth et al., 2018;Saunders et al., 2012). Enfin, par rapport aux signaux de surveillance, il a été montré pour quelques éruptions que les échelles de temps pétrologiques correspondent généralement bien à l'activité sismique ayant eu lieu des mois à quelques jours avant l'éruption (Albert et al., 2019;Longpre et al., 2014;[START_REF] Marti | Correlation of Magma Evolution and Geophysical Monitoring during the 2011-2012 El Hierro (Canary Islands) Submarine Eruption[END_REF]Pankhurst et al., 2018;[START_REF] Rasmussen | When does eruption run-up begin? Multidisciplinary insight from the 1999 eruption of Shishaldin volcano[END_REF]Ruth et al., 2018). Ces corrélations sont importantes en termes de mitigation des risques volcaniques, afin d'améliorer la gestion de futures crises volcaniques.

Ces échelles de temps des processus magmatiques sont donc longues par rapport à la durée d'une éruption, mettant quelques secondes à quelques minutes ou pouvant connaître plusieurs phases éruptives pendant plusieurs années (Cashman et al., 2017). Les dynamiques des éruptions volcaniques sont différentes en fonction de la composition du magma et de sa rhéologie, ainsi qu'en fonction de son contenu en gaz volatils et de leur comportement pendant l'ascension du magma. Cela peut conduire à des éruptions effusives ou explosives [START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF]. Les éruptions très explosives impliquent généralement un magma plus différencié, plus riche en silice donc plus visqueux et avec des gaz volcaniques piégés. Cependant, des transitions entre les différents styles éruptifs peuvent se produire durant l'histoire éruptive d'un volcan ou au cours d'une même éruption. L'étude combinée des temps de stockage magmatique, de leur corrélation ou non avec des signaux pouvant être enregistrés en surface par des observatoires volcanologiques, ainsi que la connaissance des conditions de stockage des magmas, sont donc primordiales en vue de la prévention des risques volcaniques.

Enfin, les différents objectifs de cette thèse et l'organisation du manuscrit sont présentés : les principaux objectifs sont de mieux comprendre le système de stockage magmatique des volcans dans le concept du mush et de déterminer les échelles de temps des processus pré-éruptifs, avant des éruptions monitorées ou non. Comme la prédiction d'une éruption future est impossible, nous estimons ces temps pour des éruptions anciennes afin d'en dégager des temps caractéristiques pour un système volcanique donné. Pour les éruptions monitorées dans cette thèse, nous menons une étude de cas sur les volcans du Kamchatka comme analogues de la zone de subduction des Antilles. Nous verrons si les échelles de temps pétrologiques pré-éruptives peuvent être corrélées aux signaux de surveillance et nous nous intéresserons également au phénomène de mélange des magmas via une étude expérimentale.

In this first chapter, we present a general overview on volcanoes and magma genesis as well as the changes in storage conditions that magmas can experience during their pre-eruptive history.

A volcano is an opening through the crust from which magmas erupt and hot gases can escape via a series of fractures within and beneath the volcano. When lava accumulates over long period of times, it can build a volcanic edifice that can be destabilized with time during the volcano history. Eruptions can last a short period of time (from days to months) or longer (several years). On Earth, approximately 1550 on-land volcanoes have been active in the last 10 000 years (the Holocene), with more than 9440 eruptions and approximately 800 million people that live within 100 km from a volcano [START_REF] Pankhurst | [END_REF]United Nations, 2015). By calculating the ratio of the population living in less than 30 km from an active volcano compared to its total population, 95 % of the population exposed to volcanic hazard is concentrated in six countries: Indonesia, the Philippines, Japan, Mexico, Ethiopia, Guatemala and Italy (United Nations, 2015). Volcanic hazards can have impacts on different scales and consist of primary hazards (occurring immediately during an eruption): pyroclastic density currents, ash and volcanic fragments (tephra); ballistics due to volcanic explosions; lava flows; volcanic gases and aerosols; earthquakes and lightning or secondary hazards (differed in time or space): lahars, which are destructive mixtures of volcanic debris and water, followed by heavy rainfall; volcanic instability with debris avalanches, landslides and tsunamis (for a few examples: [START_REF] Aiuppa | Forecasting Etna eruptions by real-time observation of volcanic gas composition[END_REF][START_REF] Allard | Steam and gas emission rate from La Soufriere volcano, Guadeloupe (Lesser Antilles): Implications for the magmatic supply during degassing unrest[END_REF][START_REF] Barberi | Chronology and dispersal characteristics of recently (last 5000 years) erupted tephra of Cotopaxi (Ecuador): implications for long-term eruptive forecasting[END_REF][START_REF] Boudon | Notice explicative de la carte géologique à 1/20.000e du Massif volcanique de la Soufrière[END_REF][START_REF] Cimarelli | A review of volcanic electrification of the atmosphere and volcanic lightning[END_REF][START_REF] Hildenbrand | Large-scale mass wasting on small volcanic islands revealed by the study of Flores Island (Azores)[END_REF][START_REF] Komorowski | Challenges of Volcanic Crises on Small Islands States[END_REF]Lindsay et al., 2005;Passarelli and Brodsky, 2012;United Nations, 2015). In the historical time, five major eruptions occurred: Laki in Iceland (1783, with 9,000 casualties and climate perturbations), Tambora in Indonesia (1815; 60,000 fatalities and climate perturbations), Krakatau in Indonesia (1883 with 36,000 deaths and a tsunami), Montagne Pelée in Martinique (1902;32,000 deaths) and Nevado del Ruiz in Colombia (1985;23,187 fatalities) (United Nations, 2015). The recent and massive eruption of the Hunga Tonga-Hunga Ha'apai volcano in January 2022 was the largest since Mount Pinatubo in the Philippines in 1991 (ash column of 40 km high; global cooling of 1.5 °C due to volcanic aerosols; 400 fatalities [START_REF] Newhall | Volcanic Risk Mitigation that Could Have Been Derailed but Wasn't: Pinatubo, Philippines 1991[END_REF]) and the biggest explosion ever recorded by instruments [START_REF] Cassidy | Huge volcanic eruptions: time to prepare[END_REF]. Over the next century, large-scale volcanic eruptions are hundreds of times more likely to occur than an asteroid and comet impacts put together [START_REF] Cassidy | Huge volcanic eruptions: time to prepare[END_REF][START_REF] Trilling | The Size Distribution of Near-Earth Objects Larger Than 10 m[END_REF]. Silicic magmatic systems can produce modest-to-large volumes of magmas (< 1-10 3 km 3 ) via explosive eruptions that can be a threat to populated areas, infrastructures and agriculture [START_REF] Andersen | Petrochronologic perspective on rhyolite volcano unrest at Laguna del Maule, Chile[END_REF][START_REF] Castro | Rapid ascent of rhyolitic magma at Chaitén volcano, Chile[END_REF][START_REF] Self | Consequences of Explosive Supereruptions[END_REF][START_REF] Sparks | Super-Eruptions: Global Effects and Future Threats[END_REF]. However, the world is not prepared for eruptions of high magnitudes and the probability of an eruption of a magnitude 7 eruption or greater in this century is 1 in 6 [START_REF] Lin | Magnitude, frequency and climate forcing of global volcanism during the last glacial period as seen in Greenland and Antarctic ice cores (60-9 ka)[END_REF], which caused, in the past, climate change and the collapse of civilizations [START_REF] Cassidy | Huge volcanic eruptions: time to prepare[END_REF]. Magma dynamics prior to volcanic eruptions are then important processes to constrain and understand, as the past volcanic behavior and pre-eruptive processes through the eruptive record of a volcano are a precious guide to its future behavior (Pankhurst et al., 2018). These pre-eruptive processes on specific volcanoes, and how their real-time detection in the field prior to eruptions can be assessed, are key issues in volcanology and volcanic monitoring (Pankhurst et al., 2018).

I.1. Magmas and volcanic eruptions I.1.1. Magma origin, evolution and mush model

Most active volcanoes lie at plate boundaries, where the crust is created (mid-ocean ridges, rifts) or recycled (subduction), or away from the plate boundaries, generating intraplate volcanism. Magmas can be produced following partial melting of the peridotites of the mantle but can present mineralogic and geochemical differences according to the geological context (mid-ocean ridges, intraplate volcanism, continental rifts…) or come from the partial melting of the overlying mantle, in subduction zones or oceanic crust or even sediments of the plunging plate. Different partial melting rates and depths are associated to these contexts, with ~20% of partial melting in mid-ocean ridges at 20-50 km depth, 5-15% for hot spots at 75-100 km and ~10% of partial melting in subduction zones, at 100-150 km depth [START_REF] Peycru | Géologie tout-en-un BCPST 1re et 2e années[END_REF]. By density difference compared to the host rocks, magmas rise in the crust and are stored at depths where they are neutrally buoyant for a more or less long time. In this thesis, the context in which we are interested in is subduction and it is presented in Chapter II.

Magmas are partially molten rocks from the mantle or from the crust and are constituted by crystals, silicate melt and volatile gases; their movements are modulated over long timescales by slow melt generation and transfer from the lower crust (Cashman et al., 2017). Crystals consist of periodically repeating arrays of atoms with compositions that can be represented by definite chemical formulas (Griffen, 2018). The crystals in which we are interested in are phenocrysts, which are crystals growing in the same magma where it is observed (Pankhurst et al., 2018). The two other types of crystals found in magmas are either antecrysts, which are crystals that grew in a previous magmatic system but that are recycled in another magma batch, or xenocrysts, that are accidentally captured by a new magma but which are unrelated to it [START_REF] Davidson | Microsampling and Isotopic Analysis of Igneous Rocks: Implications for the Study of Magmatic Systems[END_REF]Pankhurst et al., 2018). The evolution of magmas occurs via fractional crystallization, contamination with the host rock or mixing with other magmas. As minerals are stable at different pressure and temperature conditions, the crystallization of phenocrysts in a magma follows the Bowen reaction series, which are reactions that explain the crystallization sequence for igneous rocks (Bowen, 1928). With decreasing temperature, fractional crystallization occurs, with compatible elements (more easily incorporated in crystal lattices, as magnesium) that are incorporated first within crystals: over time, compatible elements are then withdrawn from the melt. Incompatible elements (incompatible with crystal lattices, such as alkaline elements) stay preferably in the melts. As silica tends to have the same behavior as incompatible elements, its proportion increases in the melt with time, so the melt evolves towards richer proportions in incompatible elements and silica. For Fe-Mg minerals, they crystallize in this order: olivine → pyroxene → amphibole (if water is present) → biotite mica. If the magma has intermediate-silica composition, plagioclase feldspar can appear and they evolve from calcium-rich compositions (anorthite) to more sodium-rich (albite) (Bowen, 1928). Quartz, as well as K-feldspars and muscovite mica can also be produced. Following these reactions, chemical elements are exchanged and the initial mafic liquid (45-52 % silica) evolves in a residual liquid towards more felsic compositions (> 66% silica), allowing new crystals to be formed (Bowen, 1928).

Volatile gases concentrations can vary depending on the magmas origins and their behavior during magma storage in magmatic reservoirs or during ascent before an eruption (from a few ppm to wt%) [START_REF] Balcone-Boissard | Behavior of halogens during the degassing of felsic magmas[END_REF]. H2O is the principal volatile gas in magmas but other gases are present such as CO2, SO2 and halogens (F, Cl, Br and I). Volatiles are often concentrated in the melt during fractional crystallization, as they are mostly incompatible with the solid and can reach high contents in the most evolved magmas. As their mineral/melt partition coefficients are low and their H2O vapormelt partition coefficients are high, volatiles behave simply at first order during magmatic differentiation and degassing [START_REF] Balcone-Boissard | Behavior of halogens during the degassing of felsic magmas[END_REF].

These magmas are stored in what was initially viewed as "magma chambers", where the temperature is sufficient to produce some melts but where pressure is low enough to allow the magma to exist as a liquid [START_REF] Marghany | An introduction to minerals, rocks, and mineral deposits, in: Advanced Algorithms for Mineral and Hydrocarbon Exploration Using Synthetic Aperture Radar[END_REF]. In volcanic subduction arcs, magmas are stored between 2-10 km or throughout the crust, up to the Moho, in large reservoirs or by underplating in mushy zones (Bachmann and Bergantz, 2006;[START_REF] Bergantz | Open-system dynamics and mixing in magma mushes[END_REF]Cashman and Blundy, 2013;Cashman et al., 2017;[START_REF] Christopher | Crustal-scale degassing due to magma system destabilization and magma-gas decoupling at Soufrière Hills Volcano, Montserrat. Geochemistry[END_REF]Couch et al., 2001;[START_REF] Edmonds | Extensive, water-rich magma reservoir beneath southern Montserrat[END_REF]Pinkerton et al., 2002a;[START_REF] Plank | Why do mafic arc magmas contain ∼4wt% water on average? Earth Planet[END_REF]Ruprecht et al., 2012).

The concept of a magma chamber dates back to over 100 years. Evidence has shown the existence of deep-seated magma bodies, through the presence of calderas, which are large craters that can measure several kilometers in diameter resulting from eruptions emptying a magma chamber (Cashman and Schematic of a transcrustal magmatic mush system, where melt in the mush domain results in the formation of multiple magma chambers in the middle and upper crust (modified from Cashman et al. (2017)

and Sparks et al. (2019)).

A magmatic system is then defined as four interacting domains: magma, mush, super-solidus rocks and sub-solidus host rocks while a magma reservoir is the domain within the magmatic system containing melt, fluids and crystals and is above the solidus [START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF]. Mushes are not eruptible, as their rheology is controlled by deformation of the crystalline framework (Cashman et al., 2017) and melts can be distributed on a microscale in a crystal framework (Bachmann and Huber, 2016;[START_REF] Parmigiani | Pore-scale mass and reactant transport in multiphase porous media flows[END_REF] or on a larger scale, in the form of veins, lenses or pockets that can be rapidly extracted from these zones [START_REF] Cashman | Calderas and magma reservoirs[END_REF][START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF][START_REF] Swallow | Evacuation of multiple magma bodies and the onset of caldera collapse in a supereruption, captured in glass and mineral compositions[END_REF][START_REF] Tarasewicz | Magma mobilization by downward-propagating decompression of the Eyjafjallajökull volcanic plumbing system[END_REF]. Simulations that account for particles suggest that mushes with ≤ 60% crystals can be mobilized by injection, but welded crystals or recrystallization of crystal contacts are neglected in these simulations [START_REF] Carrara | Lubrication effects on magmatic mush dynamics[END_REF][START_REF] Jarvis | Crystal and Volatile Controls on the Mixing and Mingling of Magmas[END_REF]. Furthermore, geophysical observations suggest that mushes spend the majority of their lifetimes with much higher crystallinities (80-90%) [START_REF] Farrell | Tomography from 26 years of seismicity revealing that the spatial extent of the Yellowstone crustal magma reservoir extends well beyond the Yellowstone caldera[END_REF][START_REF] Jarvis | Crystal and Volatile Controls on the Mixing and Mingling of Magmas[END_REF][START_REF] Pritchard | Synthesis: PLUTONS: Investigating the relationship between pluton growth and volcanism in the Central Andes[END_REF][START_REF] Sinton | Mid-ocean ridge magma chambers[END_REF].

Transitions in rocks can occur, when the melt is sufficient to form a mush (reaching the "melt connectivity", the threshold melt fraction being around 1-2% for basalts [START_REF] Faul | Permeability of partially molten upper mantle rocks from experiments and percolation theory[END_REF]), and when the melt and volatile content increases, it can produce a transition from mush to magma, changing the viscosity from several orders of magnitude [START_REF] Caricchi | Non-Newtonian rheology of crystal-bearing magmas and implications for magma ascent dynamics[END_REF][START_REF] Costa | A model for the rheology of particle-bearing suspensions and partially molten rocks[END_REF]. Indeed, rheology rules the magma dynamics as viscosity varies from 25 orders of magnitude in magmatic systems [START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF]. It goes from the low viscosity of gas to the high viscosity of the rocks surrounding magmatic reservoirs, playing a fundamental role in the dynamics of these systems (silicate melts: 10 0 to 10 7 Pa s compared to magmatic fluids: 10 -2 to 10 -5 Pa s) [START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF].

I.1.2. Magma dynamics

In the mush model, it appears that igneous processes vary between the upper and lower crust: with thermal and rheological conditions in the lower crust that lead to rapid melt segregation and fractionation while in the upper crust, melt-dominated magma bodies need high rates of melt transfer to maintain them, with cooler non-eruptible areas (Cashman et al., 2017).

Regarding the formation of magma chambers and reservoirs, three main phenomena can be highlighted. First, repeated magmatic intrusions up to the point of the creation of a mush zone, with sufficient growth rate and temperature can allow the formation of a magma chamber, if the melt fraction increases [START_REF] Annen | Construction and evolution of igneous bodies: Towards an integrated perspective of crustal magmatism[END_REF]. Second of all, remobilization or "defrosting" can happen, when immobile domains of a magmatic system (mush, super-solidus rock and host rock) are transformed in eruptible magma (Huber et al., 2010;[START_REF] Parmigiani | Mush microphysics and the reactivation of crystal-rich magma reservoirs[END_REF][START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF]. Remobilization can occur after a heat transfer or fluid fluxing (such as water), or both, with an injection of hydrous magma below the mush or host rock, that can lead to mixing of magma with mush [START_REF] Burgisser | A rapid mechanism to remobilize and homogenize highly crystalline magma bodies[END_REF][START_REF] Huppert | The Generation of Granitic Magmas by Intrusion of Basalt into Continental Crust[END_REF][START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF]. Finally, magma chamber formation by reactive melt segregation in the mush can take place, extracting melt from the mush to form melt-rich layers [START_REF] Bachmann | On the Origin of Crystal-poor Rhyolites: Extracted from Batholithic Crystal Mushes[END_REF][START_REF] Mckenzie | The extraction of magma from the crust and mantle[END_REF][START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF].

Magma transport through magma reservoirs occur as brittle mechanisms (sills, horizontal intrusions; dikes, vertical intrusions or laccoliths; dome-shaped), which are supported by geophysical observations, such as prior to the eruption of Tolbachik in Kamchatka in 1975-1976[START_REF] Fedotov | Seismological studies on the mechanism of the large Tolbachik fissure eruption, 1975-1976[END_REF] or geological and petrological evidence. They can also form by ductile mechanisms (diapirism) (Ruprecht and Plank, 2013;[START_REF] Scaillet | The Gangotri granite (Garhwal Himalaya): Laccolithic emplacement in an extending collisional belt[END_REF][START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF][START_REF] White | Dynamics of dyke intrusion in the mid-crust of Iceland[END_REF]. Transport of magmas within the mush also occurs, with the emplacement of new layers of magmas within a mush or at its base (called "mush replenishment"), or also by the formation of melt-rich layers within a mush due to compaction or melt flow (Figure I The merge of multiple magma and fluid layers within mush systems can happen rapidly, in only a few decades or less and could explain the very large silicic magma chamber that are associated with calderaforming eruptions (Cashman et al., 2017;[START_REF] Cashman | Calderas and magma reservoirs[END_REF]Druitt et al., 2012;Flaherty et al., 2018;[START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF]. At Yellowstone caldera (USA), tomographic studies show the presence of two large mushy zones with low-velocities bodies at 5-to 17-km depth and a deeper one (with volumes of ~10,0000 and 46,000 km 3 ), which could lead to a supereruption (Cashman et al., 2017;[START_REF] Farrell | Tomography from 26 years of seismicity revealing that the spatial extent of the Yellowstone crustal magma reservoir extends well beyond the Yellowstone caldera[END_REF][START_REF] Huang | The Yellowstone magmatic system from the mantle plume to the upper crust[END_REF]. Mush development and multiple magma layer formation is a slow process (that can occur during periods of volcano dormancy) while instability and magma transport within a mush is faster, leading to periods of volcanic unrest which could be another way of triggering eruptions [START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF]. With the upward transport of magma, it leads to a decompression linked to the exsolution of volatiles, with or without crystallization for water-saturated melts [START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF][START_REF] Tait | Pressure, gas content and eruption periodicity of a shallow, crystallising magma chamber[END_REF].

Other magma dynamics, more localized than mush reactivation and apart from magma transport are magma injection (for a few examples: Allan et al., 2013;Barker et al., 2016;Chamberlain et al., 2014;[START_REF] Davydova | Petrological evidence of rapid evolution of the magma plumbing system of Bezymianny volcano in Kamchatka before the December 20th, 2017 eruption[END_REF]Druitt et al., 2012;Matthews et al., 2012) that can be followed by magma mixing or mingling. Granular and fluid dynamics show that complex mixing can occur when magma is injected Copyrighted material, unauthorised distribution (Figure 7; [START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF] in the reservoir and that high velocity of magma injection rate can entrain crystals from the wall and bring crystals together from different parts of the reservoir that experienced different magmatic environments and storage conditions [START_REF] Bergantz | Open-system dynamics and mixing in magma mushes[END_REF]. Compositional zoning can reflect these magma processes and variations in physical conditions such as pressure, temperature, oxygen or water fugacities [START_REF] Anderson | Probable relations between plagioclase zoning and magma dynamics, Fuego Volcano, Guatemala[END_REF][START_REF] Giacomoni | Plagioclase as archive of magma ascent dynamics on "open conduit" volcanoes: The 2001-2006 eruptive period at[END_REF]Petrone and Mangler, 2021;[START_REF] Streck | Mineral Textures and Zoning as Evidence for Open System Processes[END_REF].

Magma mixing processes occurred prior to major eruptions, such as Vesuvius 79 AD, Krakatau 1883, Pinatubo 1991 and Eyjafjallajökull 2010, by refilling of a felsic volcanic magma chamber by a hotter more mafic magma [START_REF] Morgavi | Decrypting Magma Mixing in Igneous Systems[END_REF]. Mixing is the chemical interaction between two magmas that produces a composition intermediate between the original end-members [START_REF] Bunsen | Ueber die Processe der vulkanischen Gesteinsbildungen Islands[END_REF][START_REF] Jarvis | Crystal and Volatile Controls on the Mixing and Mingling of Magmas[END_REF][START_REF] Morgavi | Decrypting Magma Mixing in Igneous Systems[END_REF]. By contrast, mingling is the physical interaction of the two magmas, without chemical exchanges between them [START_REF] Jarvis | Crystal and Volatile Controls on the Mixing and Mingling of Magmas[END_REF]Morgavi et al., 2013a[START_REF] Morgavi | Decrypting Magma Mixing in Igneous Systems[END_REF][START_REF] Perugini | Analysis and numerical simulation of chaotic advection and chemical diffusion during magma mixing: petrological implications[END_REF].

Major geological evidence of magma mixing from the field can be cited extensively in many works in the last decades, as textural heterogeneity in igneous rocks [START_REF] Anderson | Magma mixing: petrological process and volcanological tool[END_REF]Auer et al., 2018;[START_REF] Bacon | Magmatic inclusions in silicic and intermediate volcanic rocks[END_REF][START_REF] De Rosa | Fractal analysis of mingled/mixed magmas: an example from the Upper Pollara eruption (Salina Island, southern Tyrrhenian Sea, Italy)[END_REF][START_REF] Didier | Enclaves and Granite Petrology[END_REF][START_REF] Eichelberger | Origin of andesite and dacite: Evidence of mixing at Glass Mountain in California and at other circum-Pacific volcanoes[END_REF][START_REF] Morgavi | Decrypting Magma Mixing in Igneous Systems[END_REF][START_REF] Morgavi | The Grizzly Lake complex (Yellowstone Volcano, USA): Mixing between basalt and rhyolite unraveled by microanalysis and X-ray microtomography[END_REF][START_REF] Perugini | Insights into magma chamber processes from the analysis of size distribution of enclaves in lava flows: A case study from Vulcano Island (Southern Italy)[END_REF][START_REF] Perugini | Analysis and simulation of magma mixing processes in 3D[END_REF]Perugini andPoli, 2012, 2005;[START_REF] Pritchard | Eruption-triggered mixing of extracaldera basalt and rhyolite complexes along the East Gallatin-Washburn fault zone[END_REF][START_REF] Rosa | Relationships between deformation and mixing processes in lava flows: a case study from Salina (Aeolian Islands, Tyrrhenian Sea)[END_REF][START_REF] Ventura | Kinematic significance of mingling-rolling structures in lava flows: a case study from Porri Volcano (Salina, Southern Tyrrhenian Sea)[END_REF][START_REF] Wada | Fractal structure of heterogeneous ejecta from the Me-akan volcano, eastern Hokkaido, Japan: implications for mixing mechanism in a volcanic conduit[END_REF], with magmatic enclaves being one evidence of this process (discrete portions of a magma present in a host magma of a different composition) [START_REF] Bacon | Magmatic inclusions in silicic and intermediate volcanic rocks[END_REF][START_REF] Davydova | Petrology of mafic enclaves in the 2006-2012 eruptive products of Bezymianny Volcano, Kamchatka[END_REF][START_REF] Didier | Enclaves and Granite Petrology[END_REF][START_REF] Walker | Some associations of acid and basic igneous rocks[END_REF][START_REF] Wilcox | Rhyolite-basalt complex on Gardiner River, Yellowstone Park, Wyoming[END_REF], compositionallybanded volcanic rocks [START_REF] Andrews | Thermal and rheological controls on the formation of mafic enclaves or banded pumice[END_REF]Auer et al., 2018;[START_REF] Clynne | A Complex Magma Mixing Origin for Rocks Erupted in 1915, Lassen Peak, California[END_REF] or disequilibrium textures in minerals (Chamberlain et al., 2014;Costa and Chakraborty, 2004;[START_REF] Martin | Bang! Month-Scale Eruption Triggering at Santorini Volcano[END_REF][START_REF] Perugini | Chaotic advection, fractals and diffusion during mixing of magmas: evidence from lava flows[END_REF]. Numerical and experimental studies of magma mixing dynamics show that chaotic mixing dynamics produce a complexity of geochemical variations and textures in the rocks [START_REF] De Campos | Decoupled convection cells from mixing experiments with alkaline melts from Phlegrean Fields[END_REF], 2008, 2011;[START_REF] Laumonier | Experimental simulation of magma mixing at high pressure[END_REF]Montagna et al., 2015;Morgavi et al., 2015;[START_REF] Perugini | Chaotic advection, fractals and diffusion during mixing of magmas: evidence from lava flows[END_REF][START_REF] Perugini | Trace element mobility during magma mixing: Preliminary experimental results[END_REF]Perugini et al., , 2015a;;[START_REF] Petrelli | Transition to chaos and implications for time-scales of magma hybridization during mixing processes in magma chambers[END_REF] and that a diffusive fractionation process occurs on chemical elements [START_REF] Perugini | Trace element mobility during magma mixing: Preliminary experimental results[END_REF][START_REF] Perugini | Diffusive fractionation of trace elements by chaotic mixing of magmas[END_REF]. Record of magma mixing processes can consist in chronometers to estimate the time between mixing and eruption, that could allow important estimates for risk assessment on several volcanoes in the world (Morgavi et al., 2013b(Morgavi et al., , 2013a[START_REF] Morgavi | Decrypting Magma Mixing in Igneous Systems[END_REF]Perugini et al., 2010Perugini et al., , 2015a;;Perugini and Poli, 2012). Numerical models show that mingling is effective on timescales of the orders of a few hours (Montagna et al., 2015;Morgavi et al., 2017;Perugini et al., 2010Perugini et al., , 2015a)).

Storage conditions can be inferred from the compositions and textures of the erupted materials (crystals, melt inclusions; magma droplets trapped in crystals or matrix phases of the transporting magma, becoming glasses) (Cashman et al., 2017;[START_REF] Moore | The hydrous phase equilibria (to 3 kbar) of an andesite and basaltic andesite from western Mexico: constraints on water content and conditions of phenocryst growth[END_REF][START_REF] Rutherford | Neodymium diffusion in orthopyroxene: Experimental studies and applications to geological and planetary problems[END_REF] , as magma storage conditions cannot be estimated in-situ (Figure I.3). Melt inclusions can give information on the original melt composition and can allow to infer the pre-eruptive magma conditions, such as temperature and pressure (and thus the depths of the magmas), magmatic volatile budgets, as well as degassing, magma ascent, partitioning behavior of elements and magmatic system evolution (Balcone-Boissard et al., 2018;[START_REF]Evidence for an Active, Transcrustal Magma System in the Last 60 ka and Eruptive Degassing Budget (H 2 O, CO 2[END_REF][START_REF] Rose-Koga | Silicate melt inclusions in the new millennium: A review of recommended practices for preparation, analysis, and data presentation[END_REF][START_REF] Schmitt | Gas-saturated crystallization and degassing in large-volume, crystal-rich dacitic magmas from the Altiplano-Puna, northern Chile[END_REF][START_REF] Wallace | Gradients in H2O, CO2, and exsolved gas in a largevolume silicic magma system: Interpreting the record preserved in melt inclusions from the Bishop Tuff[END_REF]. Matrix glass or residual glass is not trapped in the crystals and is decompressed during ascent and eruption (Figure I.3). From the residual glass and crystal rim composition, the equilibrium temperature can be inferred (Putirka, 2008). Volatiles are continuously degassed during the eruption, allowing the nucleation of bubbles and formation of a fluid phase: low volatiles content in groundmass glass with degassing of CO2, H2O, S and halogens [START_REF] Cashman | Volatile controls on magma ascent and eruption[END_REF]. By studying the textures of the residual glasses (vesicles, microlites; vesicle size distribution and crystal size distribution (CSD)), the crystallization processes can be understood, as well as magma mixing timescales, along with the processes occurring during magma ascent, as they are different depending on the ascent rates, with the rates of crystal nucleation that depend on the degree of undercooling (temperature difference between the liquidus and the magma) [START_REF] Martel | Eruption Dynamics Inferred from Microlite Crystallization Experiments: Application to Plinian and Dome-forming Eruptions of Mt. Pelee (Martinique, Lesser Antilles)[END_REF][START_REF] Mastin | Insights into volcanic conduit flow from an open-source numerical model[END_REF]. The ascent rates influence the eruptive behavior [START_REF] Cassidy | Controls on explosive-effusive volcanic eruption styles[END_REF] and depend on the pressure in the storage region, physical magma 2015; [START_REF] Lee | How important is the role of crystal fractionation in making intermediate magmas? Insights from Zr and P systematics[END_REF][START_REF] Martel | Syn-Eruptive Conditions of the AD 1530 Sub-Plinian Eruption of La Soufrière of Guadeloupe (Lesser Antilles)[END_REF][START_REF] Turner | The global chemical systematics of arc front stratovolcanoes: Evaluating the role of crustal processes[END_REF]. Crystals in the crystal cargo in the erupted rocks can record complex crystallization histories via compositional zoning, which allow to study the different residence pathways and transport in the storage regions through their zonations (Cashman and Blundy, 2013;Cashman et al., 2017;Costa et al., 2020). However, these paths can correspond to different pre-eruptive histories if crystals from different parts of the mush system were assembled shortly before the eruption [START_REF] Andrews | Repeated recharge, assimilation, and hybridization in magmas erupted from El Chichón as recorded by plagioclase and amphibole phenocrysts[END_REF][START_REF] Berlo | Textural and chemical variation in plagioclase phenocrysts from the 1980 eruptions of Mount St. Helens, USA[END_REF]Cashman and Blundy, 2013;Costa et al., 2020;Ginibre et al., 2002b;[START_REF] Kilgour | Timescales of magmatic processes at Ruapehu volcano from diffusion chronometry and their comparison to monitoring data[END_REF][START_REF] Moore | Primitive melt replenishment and crystal-mush disaggregation in the weeks preceding the 2005-2006 eruption[END_REF]Pankhurst et al., 2018;[START_REF] Smith | A Temporal Record of Magma Accumulation and Evolution beneath Nevado de Toluca, Mexico, Preserved in Plagioclase Phenocrysts[END_REF][START_REF] Wallace | Constraints on mingling of crystal populations from off-center zoning profiles: A statistical approach[END_REF], as variations in isotopic compositions in individual crystals or between them are observed [START_REF] Bindeman | Low-δ 18 O Rhyolites from Yellowstone: Magmatic Evolution Based on Analyses of Zircons and Individual Phenocrysts[END_REF]Cashman et al., 2017;[START_REF] Charlier | The Upper Crustal Evolution of a Large Silicic Magma Body: Evidence from Crystal-scale Rb-Sr Isotopic Heterogeneities in the Fish Canyon Magmatic System[END_REF][START_REF] Dungan | Partial assimilative recycling of the mafic plutonic roots of arc volcanoes: An example from the Chilean Andes[END_REF][START_REF] Lange | Diverse Sr isotope signatures preserved in mid-oceanic-ridge basalt plagioclase[END_REF].

Thus, compositional heterogeneities produced by magma mixing or disequilibrium textures in zoned crystals, frozen in time after an eruption, can be viewed as broken clocks on a crime scene, to estimate the timescales of the processes prior to an eruption [START_REF] Morgavi | Decrypting Magma Mixing in Igneous Systems[END_REF]

I.2. Timescales of magmatic processes

I.2.1. Radiochronology and crystal growth

Radiogenic dating provides the absolute age of magmatic processes (such as the age of crystallization). The study of radiogenic isotopes of the U-Th-Ra series on total rocks, separate minerals or zircons brings a greater resolution of the times of magmatic processes. Indeed, the crystallization of minerals with partition coefficients different from U-Th-Ra leads to the fractionation of elements of this series, triggering the radiogenic clocks (Costa et al., 2020). Zircons are minerals that can be recycled between individual magma batches and be used to estimate the timescales of magma differentiation (Cashman et al., 2017;[START_REF] Condomines | Timescales of Magma Chamber Processes and Dating of Young Volcanic Rocks[END_REF][START_REF] Cooper | Timescales of crustal magma reservoir processes: insights from U-series crystal ages[END_REF][START_REF] Cooper | Uranium-series Crystal Ages[END_REF]Costa et al., 2020;Morgan et al., 2004;[START_REF] Turner | Measuring Timescales of Magmatic Evolution[END_REF].

The timescales given by these crystals are of the order of 10 3 to > 10 5 years and contrast with the shorter timescales estimated for magma accumulation in the upper crust before volcanic eruptions (< 1 to 10 3 years) (Allan et al., 2013;Cashman et al., 2017;Costa et al., 2010;Druitt et al., 2012;[START_REF] Matthews | Age of the Lava Creek supereruption and magma chamber assembly at Yellowstone based on 40 Ar/ 39 Ar and U-Pb dating of sanidine and zircon crystals[END_REF]Stelten et al., 2015;[START_REF] Zellmer | Magma Emplacement and Remobilization Timescales Beneath Montserrat: Insights from Sr and Ba Zonation in Plagioclase Phenocrysts[END_REF]. However, uncertainties in the timing and U-Th-Ra series highlight the need for new methods to estimate shorter processes (Costa et al., 2020).

I.2.2. Diffusion chronometry

Diffusion chronometry can then be used to estimate these shorter timescales, using crystals as time capsules [START_REF] Chakraborty | Diffusion chronometry of volcanic rocks: looking backward and forward[END_REF]Costa, 2021;Petrone and Mangler, 2021). Diffusion of elements in a mineral occurs on much shorter timescales than radioactive isotope decay, that are unaffected by physicochemical changes in the magmatic environments, thus allowing to measure the duration of magmatic processes on shorter timescales (Chakraborty, 2008). Solid-state diffusion has many applications, such as the study of reaction mechanisms between minerals, in the field of deformation or mantle flow; weathering and maintenance of historical monuments, but also the application of solid state diffusion to determine the timescales of terrestrial or planetary processes (Chakraborty, 2008;[START_REF] Costa | Short time scales of magmatic assimilation from diffusion modeling of multiple elements in olivine[END_REF][START_REF] Faryad | Duration of Eo-Alpine metamorphic events obtained from multicomponent diffusion modeling of garnet: a case study from the Eastern Alps[END_REF]Morgan et al., 2006). Timescales of solid-state diffusion in silicates can be inferred after the eruption, as the trace of the Olivines are the 'drosophiles' of diffusion-chronometry studies, due to experimentally characterized diffusion coefficients for major, minor and some trace elements and as they are the more abundant minerals in mafic mamas [START_REF] Chakraborty | Diffusion Coefficients in Olivine, Wadsleyite and Ringwoodite[END_REF]. These studies have been carried out in active volcanic systems in Italy (Etna), Chile (Llaima), Greece (Santorini) and Hawaiʻi (Kīlauea) (Chakraborty, 2008;Costa, 2021;Costa et al., 2020;Kahl et al., 2011;Lynn et al., 2017;[START_REF] Martin | Bang! Month-Scale Eruption Triggering at Santorini Volcano[END_REF][START_REF] Rae | Time scales of magma transport and mixing at Kīlauea Volcano, Hawai'i[END_REF]Ruth et al., 2018) (Figure I.5b). In addition, orthopyroxene crystals are increasingly studied (Chamberlain et al., 2014;Davydova et al., 2018;Fabbro et al., 2018;[START_REF] Kilgour | Timescales of magmatic processes at Ruapehu volcano from diffusion chronometry and their comparison to monitoring data[END_REF][START_REF] Mangler | Interplinian effusive activity at Popocatépetl volcano, Mexico: New insights into evolution and dynamics of the plumbing system[END_REF][START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF][START_REF] Sato | Pre-eruptive magmatic processes of historical activities at Zao volcano, northeastern Japan: Insights from compositional zoning in orthopyroxene phenocrysts[END_REF]Saunders et al., 2012;Singer et al., 2016;Solaro et al., 2020)

(Figure I.5b).
Some studies have already been carried out on several systems, involving various magma compositions and eruptive dynamics as well as the use of different crystal/element pairings. Few studies use diffusion timescales from different elements within a given mineral (Albert et al., 2019[START_REF] Albert | Timing of Magmatic Processes and Unrest Associated with Mafic Historical Monogenetic Eruptions in Tenerife Island[END_REF][START_REF] Andersen | Petrochronologic perspective on rhyolite volcano unrest at Laguna del Maule, Chile[END_REF][START_REF] Costa | Short time scales of magmatic assimilation from diffusion modeling of multiple elements in olivine[END_REF]de Maisonneuve et al., 2016;[START_REF] Devoir | Residence time of igneous garnet in Si-rich magmatic systems: Insights from diffusion modeling of major and trace elements[END_REF]Druitt et al., 2012;[START_REF] Gordeychik | Growth of, and diffusion in, olivine in ultra-fast ascending basalt magmas from Shiveluch volcano[END_REF]Lynn et al., 2017;Morgan et al., 2006;[START_REF] Mutch | Millennial storage of near-Moho magma[END_REF]Ruth et al., 2018;Till et al., 2015) and fewer have modelled diffusion in multiple minerals from the same eruption [START_REF] Brugman | Common assumptions and methods yield overestimated diffusive timescales, as exemplified in a Yellowstone post-caldera lava[END_REF]Chamberlain et al., 2014;[START_REF] Cooper | Rapid assembly and rejuvenation of a large silicic magmatic system: Insights from mineral diffusive profiles in the Kidnappers and Rocky Hill deposits, New Zealand[END_REF]Fabbro et al., 2018;Flaherty et al., 2018;[START_REF] Magee | The Lead-up to Mount Etna's Most Destructive Historic Eruption[END_REF][START_REF] Shamloo | Decadal transition from quiescence to supereruption: petrologic investigation of the Lava Creek Tuff, Yellowstone Caldera[END_REF]Singer et al., 2016) or modelled more than one compositional boundary (rim) (Druitt et al., 2012;Kahl et al., 2011;Petrone et al., 2018Petrone et al., , 2016;;[START_REF] Rout | [END_REF]

(Figure I.5b).
Based on three types of crystals: sanidines (Ba-Sr diffusion), quartz (Ti) and pyroxenes (Fe-Mg) from the Bishop Tuff eruption, different timescales were found (1-10 years for orthopyroxenes, 10-100 for quartz and 100-1000 for sanidines), thus remobilizing 600 km 3 of magma on timescales ranging from several tens of years to a millennium before the eruption (Chamberlain et al., 2014). These differences can rise from the different diffusion rates of the elements, e.g. a fast diffusing element would reequilibrate following each magmatic event and would therefore record the last thermal or chemical disturbance before the eruption. Furthermore, combining orthopyroxenes and plagioclase of eruptions preceding caldera-forming eruptions of Santorini, similar residence timescales were found, from years to centuries showing magma movements in the crustal reservoir (Fabbro et al., 2018). Meaningful timescales can be retrieved from elemental diffusion stratigraphy when knowledge of the mineralogy, geochemistry (crystals and melt compositions) and storage conditions are known as well as a good statistic of modeled crystals. It has been resolved for Fe-Mg interdiffusion of Stromboli and Popocatépetl volcanoes suggesting different magma dynamics, with short residence times for Stromboli (~1 year or less) (Petrone et al., 2018) compared to several decades for Popocatépetl, where magma hybridization occurred in a long-lived and stable extensive plumbing system [START_REF] Mangler | Interplinian effusive activity at Popocatépetl volcano, Mexico: New insights into evolution and dynamics of the plumbing system[END_REF][START_REF] Mangler | Magma recharge patterns control eruption styles and magnitudes at Popocatépetl volcano (Mexico)[END_REF]Petrone and Mangler, 2021).

New approaches of diffusion modelling have been developed in the last few years, combining diffusion of several elements (Fe-Mg and Al content as initial conditions in clinopyroxenes of Campi Flegrei; [START_REF] Pelullo | Insights into the temporal evolution of magma plumbing systems from compositional zoning in clinopyroxene crystals from the Agnano-Monte Spina Plinian eruption (Campi Flegrei, Italy)[END_REF]) or multiple elements in multiple mineral phases and taking into account a general petrological approach [START_REF] Brugman | Common assumptions and methods yield overestimated diffusive timescales, as exemplified in a Yellowstone post-caldera lava[END_REF] Diffusion techniques then provide numerous time constraints on magmatic processes preceding an eruption, for elements whose diffusion coefficients are known, notably through experimental studies (Costa et al., 2020).

For unmonitored volcanoes, magma ascent rate studies can help mitigating volcanic hazards, using fast diffusing volatile elements (H, Li) in minerals or melt embayments, formed during phenocryst growth with a volatile content corresponding to the initial volatile content of the external melt (Costa et al., 2020). A chemical gradient between the ascending magma, crystal compositions and liquid pockets trapped at higher depths appears, as volatiles are exsolved in a vapour phase (Costa et al., 2020). These rates range from 0.01 to 0.1 MPa.s -1 for silicic eruptions of Yellowstone, Long Valley (USA) and Taupo, (New-Zealand), which is of an order of magnitude lower than what is estimated on bubble nucleation (recording the final moments of decompression, before magma fragmentation) (Costa et al., 2020;[START_REF] Myers | Ascent rates of rhyolitic magma at the onset of three caldera-forming eruptions[END_REF][START_REF] Myers | Prolonged ascent and episodic venting of discrete magma batches at the onset of the Huckleberry Ridge supereruption, Yellowstone[END_REF][START_REF] Toramaru | BND (bubble number density) decompression rate meter for explosive volcanic eruptions[END_REF].

The magmatic processes thus occur on relatively slow timescales, differing from volcanic eruptions where large volumes of magmas are emitted in a short period of time (Cashman et al., 2017). Timescales on several of these processes can help understand how the volcanic eruptions are triggered, as for Popocatépetl, where it was shown that effusive eruptions are preceded by shorter recharge activity compared to magma recharge prior to explosive eruptions [START_REF] Mangler | Magma recharge patterns control eruption styles and magnitudes at Popocatépetl volcano (Mexico)[END_REF]. Explosive eruptions record a higher number of individual recharge episodes, with the largest eruptions characterized by an ~1 year recharge hiatus directly prior to the eruption [START_REF] Mangler | Magma recharge patterns control eruption styles and magnitudes at Popocatépetl volcano (Mexico)[END_REF].

I.3. Volcanic eruptions and eruptive styles

Volcanic eruptions are traditionally divided between effusive or explosive behavior but transitions between eruptive styles and phases can occur in the eruptive history of a volcano or during the same eruption (Figure I.6). This thesis is focused on different eruptions, some of them having experienced multiple styles with extrusive and explosive phases.

The difference between the effusive and explosive eruptions is due to magma composition, its rheology and to the volatile content and its behaviour during the ascent of the magma, which is controlled by composition, dissolved H2O content, temperature of the melt phase and volumetric proportions and physical properties of suspended crystals and bubbles [START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF]. Explosive eruptions generally involve more differentiated magma, richer in silica and thus more viscous, with trapped volcanic gases. Eruptions dynamics are generally linked to the open or close nature of the volcanic system and volatiles content, which can result in radically different eruptive styles.

Volatiles are often exsolved during magma ascent, as volatile element solubilities decrease with decreasing pressure [START_REF] Balcone-Boissard | Behavior of halogens during the degassing of felsic magmas[END_REF]. They either leave the magmas and rise to the surface, if the system is open, or stay trapped in the volcano with the melt, if the system is closed. To measure the extent of an eruption, the Volcano Explosivity Index (VEI) (Newhall and Self, 1982) can be estimated, it is a general indicator of the explosive character of an eruption, taking into account the magnitude (from the volume of ejecta), intensity, destructiveness, dispersive power, violence [START_REF] Walker | The Taupo pumice: Product of the most powerful known (ultraplinian) eruption?[END_REF] and energy release, going from 0 to 8 (Figure I.6j). Eruptions having a VEI equal to or greater than 4 are considered as major eruptions, with the example of Tambora eruption in Indonesia ( 1815), with a VEI of 7 (Newhall and Self, 1982). Hawaiian and Strombolian eruptions are relatively small explosive eruptions (VEI 1-3 [START_REF] Pyle | Sizes of Volcanic Eruptions[END_REF]), produced by silica poor and low viscosity magmas. Hawaiian eruptions (from Kilauea, Hawaiʻi) begin with a fissure formed from dyke propagation and produce lava fountains from tens to hundreds of meters in height above the vent that can result in lava flows [START_REF] Gurioli | Dynamics of Strombolian explosions: Inferences from field and laboratory studies of erupted bombs from Stromboli volcano[END_REF][START_REF] Mangan | The dynamics of Hawaiian-style eruptions: a century of study[END_REF] (Figure I.6a). Strombolian eruptions (from Stromboli, Italy) result in short-lived bursting of gas pockets and eject pyroclasts to a few tens or hundreds of meters high [START_REF] Francis | Volcanoes. 2nd Edition[END_REF][START_REF] Houghton | Stronger or longer: Discriminating between Hawaiian and Strombolian eruption styles[END_REF][START_REF] Taddeucci | Hawaiian and Strombolian Eruptions[END_REF] (Figure I.6b). It can also involve ash explosions (~2 km) and violent Strombolian activity with increasing magma and gas supply rate (as well as bulk magma viscosity).

Vulcanian eruptions are named after Vulcano (Aeolian Islands, Italy) and generally involve crystal rich, high silica magmas and are associated with lava domes, with an impermeable plug below which magma in the conduit is crystallizing and becomes pressurized until the point of the eruption (VEI 2-3 [START_REF] Pyle | Sizes of Volcanic Eruptions[END_REF]) [START_REF] Burgisser | Conduit processes during the February 11, 2010 Vulcanian eruption of Soufrière Hills, Montserrat[END_REF][START_REF] Clarke | Vulcanian eruptions, in: Encyclopedia of Volcanoes[END_REF][START_REF] Francis | Volcanoes. 2nd Edition[END_REF]. Small magnitude (< 10 m 3 ) and high eruptive columns are produced and can reach 10-20 km [START_REF] Francis | Volcanoes. 2nd Edition[END_REF]) (Figure I.6c). These eruptions can result in the destruction of some parts of the volcano, with the possibility of producing pyroclastic density currents [START_REF] Burgisser | Conduit processes during the February 11, 2010 Vulcanian eruption of Soufrière Hills, Montserrat[END_REF][START_REF] Clarke | Vulcanian eruptions, in: Encyclopedia of Volcanoes[END_REF][START_REF] Francis | Volcanoes. 2nd Edition[END_REF]. Plinian eruptions are associated to the highest VEI (VEI > 4 [START_REF] Pyle | Sizes of Volcanic Eruptions[END_REF]) and are named after the 79 AD eruption of Vesuvius (Italy) by Pline le Jeune [START_REF] Cioni | Plinian and Subplinian Eruptions[END_REF]. They generally involve silica and volatiles rich magmas and consist of volcanic pulses with sustained convective columns due to a continuous magma alimentation, emitting tephra (up to 35-55 km and can last for hours) [START_REF] Cioni | Plinian and Subplinian Eruptions[END_REF][START_REF] Francis | Volcanoes. 2nd Edition[END_REF]. Phases of column collapse can happen and may altern with other eruptive styles, such as Vulcanian explosions whit shorter episodes or lava domes [START_REF] Cioni | Plinian and Subplinian Eruptions[END_REF][START_REF] Francis | Volcanoes. 2nd Edition[END_REF]

(Figure I.6d).
Unlike lava flows, lava can accumulate in the volcanic conduit and in the volcanic edifice, by internal push at the vent of relatively high viscous lava (andesite to rhyolite). When dense and viscous magmas arrive at the surface to form a lava dome, gas escaping from the conduit can produce collapse of the vesicles in magmas, microcrystallization and increase the viscosity (Boudon et al., 2015;[START_REF] Newhall | Explosive activity associated with the growth of volcanic domes[END_REF] (Figure I.6e). The morphologies of the dome are different: if the lava is not too viscous, it forms a dome that grows by internal feeding and vertical expansion (either with high internal cohesion or lowered lava dome, with a lower internal cohesion). High pressures of the gases in some lava domes can lead to two types of dome-forming eruptions with pyroclastic activity: collapse of unstable parts of the lava dome producing block-and ash-concentrated pyroclastic density currents (few tens of m/s) when an evolved degassed magma arrives at the surface (Boudon et al., 2015;[START_REF] Druitt | Pyroclastic density currents[END_REF] (Figure I.6g). Another type of dome-forming eruption can occur when a less degassed batch of magma arrives at the base of the lava dome and can generate a superficial explosion that is generally laterally directed, producing highly destructive block-and ash-diluted and turbulent pyroclastic density currents (hundreds of m/s or more) (Boudon et al., 2015;[START_REF] Bourdier | Stratigraphy of the 1902 and 1929 nuée-ardente deposits[END_REF][START_REF] Druitt | Pyroclastic density currents[END_REF][START_REF] Lacroix | La Montagne Pelée et ses éruptions[END_REF] (Figure I.6f). This explosive eruption type occurred at Montagne Pelée (Martinique) in 1902, when the lava dome collapsed and formed block-and ash-diluted and turbulent pyroclastic density currents that destroyed the south-western flank of the volcano and the town of Saint Pierre and responsible for the death of 32,000 inhabitants (Boudon et al., 2015;[START_REF] Bourdier | Stratigraphy of the 1902 and 1929 nuée-ardente deposits[END_REF][START_REF] Lacroix | La Montagne Pelée et ses éruptions[END_REF]

(Figure I.6f).
Composite lava domes can also be produced, when a dome grows and overflows laterally, forming a thick lava flow, that can grow for months or years, as it occurred in Bezymianny (Kamchatka) or Merapi (Indonesia), for examples [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF]. These domes have complex structures due to the succession of several construction and destruction phases. Lava domes can also be hidden in a volcanic edifice, producing a cryptodome, such as Mount St Helens or Bezymianny before its 1956 flank collapse [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF]. The explosion of Mount St Helens in 1980 resulted from the collapse of part of the flank of the volcano due to the instability related to the injection of the cryptodome that formed during the month preceding the eruption [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF][START_REF] Christiansen | Chronology of the 1980 eruptive activity[END_REF]

(Figure I.6h-i).
The sudden decompression of the cryptodome led to a laterally directed blast in 10-20 s, destroying an area of ~600 km 2 and which reached a maximum distance of 25 km from the vent [START_REF] Christiansen | Chronology of the 1980 eruptive activity[END_REF]

(Figure I.6i).
Explosive eruptions forming pyroclastic density currents will be briefly presented, as this thesis mainly focuses on the type of deposits generated by them. Pyroclastic density currents are produced from volcanic eruptions from small-volume events (< 1 km 3 ) to caldera-forming eruptions (100-1000 km 3 ) and are mixtures of fragmented magma, magmatic and entrained gases and lithic clasts derived from the volcanic vent, conduit and host-rock [START_REF] Dufek | Pyroclastic Density Currents[END_REF] (Figure I.6b-7). They are a type of particle gravity currents that propagate laterally away from the source region, as their density is superior to the ambient atmosphere due to different mechanisms (Figure I.7). They produce a wide range of deposits, depending on the temperatures: forming completely nonwelded deposits or ignimbrites that are extremely welded [START_REF] Dufek | Pyroclastic Density Currents[END_REF].

Finally, hundreds of volcanoes around the world may be active and should be monitored [START_REF] Caricchi | There are more active volcanoes than you think[END_REF][START_REF] Giordano | Determining the State of Activity of Transcrustal Magmatic Systems and Their Volcanoes[END_REF]. Recently, a study identified the volcanoes that are likely to reactivate, considering their last eruption, as well as the thermal state of their plumbing system [START_REF] Caricchi | There are more active volcanoes than you think[END_REF][START_REF] Giordano | Determining the State of Activity of Transcrustal Magmatic Systems and Their Volcanoes[END_REF]. This new Volcanic Activity Index identifies the volcanoes that need to be well-monitored (Kelud (Indonesia), Puyehue-Cordón Caulle (Chile), Aira (Japan), Rabaul (Papua New Guinea), St. Helens (USA)…) [START_REF] Caricchi | There are more active volcanoes than you think[END_REF][START_REF] Giordano | Determining the State of Activity of Transcrustal Magmatic Systems and Their Volcanoes[END_REF]. Investigation of the plumbing system of these volcanoes up to 15-20 km is needed, with new techniques that would allow to investigate deeper than what is possible with seismic tomography or electrical conductivity [START_REF] Caricchi | There are more active volcanoes than you think[END_REF], as well as their close monitoring.

I.4. Correlation of petrological signals with monitoring data

Approximately 80 % of the 1551 on-land active volcanoes are not adequately studied and even fewer are actively monitored (Pankhurst et al., 2018;United Nations, 2015). Unrest is a manifestation of complex subsurface processes which might or might not indicate an imminent eruption [START_REF] Gottsmann | Volcanic Unrest, Advances in Volcanology[END_REF]. Management of volcanic crisis are a challenge for civil protection authorities regarding the limited time to take important decisions and to activate protection plans (Utami et al., 2021). To improve the management of volcanic crisis and detect potential unrest, monitoring of volcanoes is performed through real-time acquisition and interpretation of different signals acquired by instrumental networks:

Copyrighted material, unauthorised distribution (Figure 35.1;[START_REF] Dufek | Pyroclastic Density Currents[END_REF] seismic, geodetic and gas geochemistry (Re et al., 2021;[START_REF] Sigmundsson | Intrusion triggering of the 2010 Eyjafjallajökull explosive eruption[END_REF][START_REF] Tarasewicz | Magma mobilization by downward-propagating decompression of the Eyjafjallajökull volcanic plumbing system[END_REF][START_REF] Zuccarello | The coupling between very long period seismic events, volcanic tremor, and degassing rates at Mount Etna volcano[END_REF].

A recent study also showed the potential of magnetic data that can be used to monitor volcanoes by producing similar data to gravimetric techniques for a lower cost [START_REF] Biasi | Volcano Monitoring With Magnetic Measurements: A Simulation of Eruptions at Axial Seamount, Kīlauea, Bárðarbunga, and Mount Saint Helens[END_REF]. Magnetic models for the different volcanic systems and best-practices have to be developed but it could be used in developing nations with high-volcanic risk [START_REF] Biasi | Volcano Monitoring With Magnetic Measurements: A Simulation of Eruptions at Axial Seamount, Kīlauea, Bárðarbunga, and Mount Saint Helens[END_REF].

As practically every eruption is preceded by an increase in seismic activity, it is used for eruption forecasting. The main seismic signals in a volcanic context are now briefly exposed, as a correlation between petrological and seismic signals will be studied in this thesis.

An earthquake is a sudden release of energy in the lithosphere that creates seismic waves (primary or secondary waves (P or S) (Acocella, 2021). To measure the size of an earthquake, the magnitude M and seismic moment M0 are used. M is the amplitude of the elastic waves corrected for the distance, on the basis of the Richter scale, where a change of one unit in magnitude corresponds to a change in radiated energy by a factor of 30, whereas M0 is the product of fault area, fault slip and shear modulus of the rock (Acocella, 2021). Magmatic processes can generate volcanic earthquakes prior to and during eruptions, which are located within a few tens of kilometers below volcanoes, as swarms (groups of several events with similar location and magnitude, without a main shock) (Acocella, 2021;[START_REF] Mcnutt | Volcanic seiscmity[END_REF]. There are different types of volcanic earthquakes, mainly distinguished by their frequency: highfrequency earthquakes (volcano-tectonic), low frequency (long-period) or volcanic tremor (Acocella, 2021;[START_REF] Mcnutt | Volcanic seiscmity[END_REF]

(Figure I.8a).
Volcano-tectonic earthquakes, are caused by shear failure along faults, due to pressure of intruding magma (for example, during dike propagation) [START_REF] Kilburn | Multiscale fracturing as a key to forecasting volcanic eruptions[END_REF][START_REF] Tokarev | Forecasting volcanic eruptions from seismic data[END_REF], pressure from degassing processes or fracture of the magma [START_REF] Lavallée | Seismogenic lavas and explosive eruption forecasting[END_REF][START_REF] Menand | A phenomenological model for precursor volcanic eruptions[END_REF]Passarelli and Brodsky, 2012) and mainly occur in the upper crust (Acocella, 2021;[START_REF] Journeau | Seismic tremor reveals active trans-crustal magmatic system beneath Kamchatka volcanoes[END_REF]. They often delineate geologic structures (rift zone, crustal detachment, etc.) existing in the vicinity of volcanoes rather than magma/melt-containing regions [START_REF] Journeau | Seismic tremor reveals active trans-crustal magmatic system beneath Kamchatka volcanoes[END_REF][START_REF] Matoza | Systematic relocation of seismicity on Hawaiʻi Island from 1992 to 2009 using waveform cross correlation and cluster analysis[END_REF][START_REF] Roman | The origin of volcano-tectonic earthquake swarms[END_REF]. The second type, long-period earthquakes, are thought to be generated by fluid pressurization processes, such as bubble formation and collapse [START_REF] Aki | Seismic monitoring and modeling of an active volcano for prediction[END_REF][START_REF] Benson | Laboratory Simulation of Volcano Seismicity[END_REF][START_REF] Burlini | Seismicity preceding volcanic eruptions: New experimental insights[END_REF] and deep long-period earthquakes have also been found in some volcanic settings (down to 40 km depth) and could correspond to a magmatic or fluid activity (Acocella, 2021). Ultra-long-periods earthquakes have also been identified, with recent progress in volcano monitoring and numerical simulations which show that these signals can correspond to magma injection in magma chambers beneath volcanoes [START_REF] Longo | Magma convection and mixing dynamics as a source of Ultra-Long-Period oscillations[END_REF]. Finally, volcanic tremor is a continuous signal that can last from minutes to days, or even longer and is interpreted as a sign of an impending eruption as it is inferred to be caused by magma propagating in a shallow dike, close to the surface (Acocella, 2021). Time history of a general volcanic earthquake swarm signals recorded before an eruption are presented in Figure I.8a, along with dominant magmatic processes causing the seismic signals.

The duration of seismic unrest in some way then informs on the magmatic motion at depth and its effects through the hosting crust prior to an eruption (Passarelli and Brodsky, 2012). be associated to the way eruptions were triggered, with eruptions with a rapid onset of seismicity prior to the eruption that may be triggered internally, from a magma crystallizing that cools down, generating volatile saturation and high overpressures, leading to fast magma ascent. While longer unrest times, with deformation signals, may be caused by magma injections prior to ascent and eruption [START_REF] Cassidy | Explosive Eruptions With Little Warning: Experimental Petrology and Volcano Monitoring Observations From the 2014 Eruption of Kelud, Indonesia. Geochemistry[END_REF]. Deformation was the first sign of unrest in less than half of the eruptions presented in Figure I.9b, often beginning months to years prior to the eruptions or prior to any other sign of unrest, with a short run-up of seismic activity [START_REF] Cassidy | Explosive Eruptions With Little Warning: Experimental Petrology and Volcano Monitoring Observations From the 2014 Eruption of Kelud, Indonesia[END_REF]. However, unrest signals such as deformation are not necessarily a sign of imminent eruption and some magma bodies can reside in the crust for months to decades before eruption [START_REF] Biggs | Global link between deformation and volcanic eruption quantified by satellite imagery[END_REF][START_REF] Ebmeier | Synthesis of global satellite observations of magmatic and volcanic deformation: implications for volcano monitoring & the lateral extent of magmatic domains[END_REF][START_REF] Jay | Locating magma reservoirs using InSAR and petrology before and during the 2011-2012 Cordón Caulle silicic eruption[END_REF]. Eruptions have also taken place without recording any deformation [START_REF] Cassidy | Explosive Eruptions With Little Warning: Experimental Petrology and Volcano Monitoring Observations From the 2014 Eruption of Kelud, Indonesia[END_REF][START_REF] Rasmussen | When does eruption run-up begin? Multidisciplinary insight from the 1999 eruption of Shishaldin volcano[END_REF] or with only a few hours of pre-eruptive seismicity (Arzilli et al., 2019b).

These geophysical and geochemical signals can be linked to the changes in the magmatic system with magma transfer through the crust within short run-up times, of a few hours to a few weeks preceding an eruption [START_REF] Ágústsdóttir | Intense Seismicity During the 2014-2015 Bárðarbunga-Holuhraun Rifting Event, Iceland, Reveals the Nature of Dike-Induced Earthquakes and Caldera Collapse Mechanisms[END_REF][START_REF] Aiuppa | Unusually large magmatic CO2 gas emissions prior to a basaltic paroxysm[END_REF][START_REF] Aiuppa | The 2007 eruption of Stromboli volcano: Insights from real-time measurement of the volcanic gas plume CO2/SO2 ratio[END_REF][START_REF] Aiuppa | Forecasting Etna eruptions by real-time observation of volcanic gas composition[END_REF][START_REF] Alparone | A multidisciplinary approach to detect active pathways for magma migration and eruption at Mt. Etna (Sicily, Italy) before the 2001 and 2002-2003 eruptions[END_REF][START_REF] Bonaccorso | Insight on recent Stromboli eruption inferred from terrestrial and satellite ground deformation measurements[END_REF][START_REF] Di Lieto | Strain Signals Before and During Paroxysmal Activity at Stromboli Volcano, Italy[END_REF][START_REF] Peltier | Cyclic magma storages and transfers at Piton de La Fournaise volcano (La Réunion hotspot) inferred from deformation and geochemical data[END_REF]Re et al., 2021;[START_REF] Ripepe | The onset of the 2007 Stromboli effusive eruption recorded by an integrated geophysical network[END_REF]Seropian et al., 2021;[START_REF] Tarasewicz | Magma mobilization by downward-propagating decompression of the Eyjafjallajökull volcanic plumbing system[END_REF]. However, monitoring efforts do not allow a prediction of when a volcanic eruption will occur, or their explosivity or even their evolution with time as understanding of the magmatic processes controlling the eruptions and their associated unrest patterns are poorly constrained (Utami et al., 2021). Despite that, the number of casualties per eruption has decreased in the last few decades, due to improvement of volcano monitoring, hazard assessment, early warning, forecast and communication around several volcanoes, with 50,000 lives saved over the last century [START_REF] Auker | A statistical analysis of the global historical volcanic fatalities record[END_REF]United Nations, 2015). What then may be of interest for better management of volcanic crises is to know the time remaining between changes in storage conditions and the start of the eruption, which can be done for past monitored eruptions if the eruption date is known.

Petrological monitoring can be undertaken on specific volcanoes and focus on volcanic products, aiming at understanding the history of the magma, storage zones, pathways of magma transport and timescales of magmatic processes (Albert et al., 2019;[START_REF] Berthod | The 2018-ongoing Mayotte submarine eruption: Magma migration imaged by petrological monitoring[END_REF][START_REF] Corsaro | Petrologic evidence of a complex plumbing system feeding the July-August 2001 eruption of Mt. Etna, Sicily, Italy[END_REF][START_REF] D'oriano | Ash erupted during normal activity at Stromboli (Aeolian Islands, Italy) raises questions on how the feeding system works[END_REF][START_REF] Muro | The Shallow Plumbing System of Piton de la Fournaise Volcano (La Reunion Island, Indian Ocean) Revealed by the Major 2007 Caldera-Forming Eruption[END_REF][START_REF] Kahl | Compositionally zoned crystals and real-time degassing data reveal changes in magma transfer dynamics during the 2006 summit eruptive episodes of Mt. Etna[END_REF][START_REF] Liu | Petrologic monitoring at Volcán de Fuego, Guatemala[END_REF][START_REF] Meletlidis | Xenopumices from the 2011-2012 submarine eruption of El Hierro (Canary Islands, Spain): Constraints on the plumbing system and magma ascent[END_REF][START_REF] Meletlidis | New insight into the 2011-2012 unrest and eruption of El Hierro Island (Canary Islands) based on integrated geophysical, geodetical and petrological data[END_REF]Re et al., 2021;Sundermeyer et al., 2020a;[START_REF] Taddeucci | Conduit processes during the July-August 2001 explosive activity of Mt. Etna (Italy): inferences from glass chemistry and crystal size distribution of ash particles[END_REF] (Figure I.9a). Correlating these pre-or syn-eruptive diffusion timescales with real-time signals obtained by instrumental networks has been done in several studies and in several volcanic contexts, from hot spots to arc volcanoes: Ruapehu volcano (1969,1971,1977,1995,1996) [START_REF] Kilgour | Timescales of magmatic processes at Ruapehu volcano from diffusion chronometry and their comparison to monitoring data[END_REF]; El Hierro (2011-2012) (Longpre et al., 2014;[START_REF] Marti | Correlation of Magma Evolution and Geophysical Monitoring during the 2011-2012 El Hierro (Canary Islands) Submarine Eruption[END_REF]; Eyjafjallajökull (2010) (Pankhurst et al., 2018);Bezymianny (2006Bezymianny ( -2012) ) (Davydova et al., 2018);Etna (1991Etna ( -1993 ; ;2006) [START_REF] Kahl | Compositionally zoned crystals and real-time degassing data reveal changes in magma transfer dynamics during the 2006 summit eruptive episodes of Mt. Etna[END_REF](Kahl et al., , 2011)); Piton de La Fournaise (2007 ;2014-2015) (Albert et al., 2019;Sundermeyer et al., 2020a); Kelud (1990) [START_REF] Utami | [END_REF]), Llaima (2008) (Ruth et al., 2018) (Figure I.9a); Shishaldin (1999) [START_REF] Rasmussen | When does eruption run-up begin? Multidisciplinary insight from the 1999 eruption of Shishaldin volcano[END_REF]Mt St Helens (1980-1986) (Saunders et al., 2012); Kīlauea (1959) [START_REF] Rae | Time scales of magma transport and mixing at Kīlauea Volcano, Hawai'i[END_REF]Merapi (2006 ;[START_REF] Chakraborty | Diffusion Coefficients in Olivine, Wadsleyite and Ringwoodite[END_REF] (Costa et al., 2013);Irazú (1963Irazú ( -1965) ) (Ruprecht and Plank, 2013). Correlation between geodetic signals (deformation, InSAR) and petrological information on the evolution of the magma bodies has also been done (Caricchi et al., 2014b;[START_REF] Jay | Locating magma reservoirs using InSAR and petrology before and during the 2011-2012 Cordón Caulle silicic eruption[END_REF], as well as linking seismicity to fluid interactions and magma migration via laboratory experiments or simulations [START_REF] Benson | Laboratory Simulation of Volcano Seismicity[END_REF][START_REF] Burlini | Seismicity preceding volcanic eruptions: New experimental insights[END_REF][START_REF] Roman | Seismological evidence for long-term and rapidly accelerating magma pressurization preceding the 2009 eruption of Redoubt Volcano, Alaska[END_REF] or thermobarometric constraints to seismicity and ground deformation [START_REF] Chiarabba | Structural features of the shallow plumbing system of Vulcano Island Italy[END_REF][START_REF] Patanè | Magma Ascent and the Pressurization of Mount Etna's Volcanic System[END_REF][START_REF] Widiyantoro | Seismic imaging and petrology explain highly explosive eruptions of Merapi Volcano[END_REF]. Short timescales are found, from the orders of hours to years before the eruptions, which are generally associated to magma injection or magma mixing processes prior to the eruption (Figure I.9a). For the 2010 eruption of Eyjafjallajökull, olivine crystals provided evidence of shallow magma intrusion and earthquakes that overlapped with GPS stations displacement data or magma migration was shown to have occurred at shallow levels six months before the 2008 eruption of Llaima volcano (Chile) (Pankhurst et al., 2018;Ruth et al., 2018)

I.5. Aims and organization of the manuscript

As seen throughout this chapter, volcanoes and arc volcanism produce destructive eruptions, generally associated to numerous casualties and devastation, leading to major societal and economical effects. Improving our knowledge on the time remaining between the first unrest signals and the eruption on specific volcanoes is then fundamental. The correlation between monitoring signals and petrological timescales is important, especially for volcanoes whose last magmatic eruptions have not been monitored, as is the case for La Soufrière of Guadeloupe or Montagne Pelée of Martinique (Lesser Antilles). As volcanic eruptions are not predictable, knowing this time on past eruptions of these volcanoes by looking at crystals carried by the magma during an eruption and ensuring a good correlation with precursory signals for monitored eruptions is important data for the management of a future volcanic crisis on one of these volcanoes.

We then aim to answer to these key issues in volcanology:

-What are the dynamics experienced by magma plumbing systems of volcanoes in the concept of mush?

-What are the timescales of the pre-eruptive processes, prior to monitored or unmonitored eruptions?

-For monitored eruptions, can these pre-eruptive petrological timescales be correlated to the monitoring signals?

-How do mixing and mingling occur within the framework of crystal mushes?

My thesis is part of the V-Care (VolcaniC early awAREness) ANR project of my supervisors, which aim to determine new indicators for a better management of volcanic crises and in particular, an early warning system for dangerous volcanoes in subduction zones, based on petrological and geophysical investigations. One of the objectives is also to estimate the time between readjustments in the magma mush and the onset of eruption on past eruptions. Indeed, first studies using diffusion timescales modelling in orthopyroxenes from several eruptions of Montagne Pelée (Martinique) in the Lesser Antilles arc showed a systematic pattern in the timescales modelled between changes in magma storage conditions and eruptions (for the 5 past Plinian-type explosive eruptions in the last 5000 years) (Boudon et al., 2018).

This systematic approach over time, although with different disruptive processes, is of great interest in terms of crisis management for a future eruption on a volcano that is certainly monitored but that has never erupted since the beginning of volcanological monitoring. Little work has been done on these correlations for volcanoes in active subduction zones with differentiated magmas (Davydova et al., 2018;[START_REF] Kilgour | Timescales of magmatic processes at Ruapehu volcano from diffusion chronometry and their comparison to monitoring data[END_REF]Saunders et al., 2012;Utami et al., 2021). However, they are fundamental in terms of volcanic risk and crisis management so that the population and administrative authorities can take the necessary measures.

It is therefore interesting to look at well-monitored recent eruptions to see if the correlation between the readjustments in the reservoir and the precursory signals is shown. If the storage area is stable over time (same magma depth, same stored volumes, same H2O saturation conditions...), this will provide an advancement towards early warning clocks in the event of a reactivation of these volcanoes.

Also, during my last year of PhD, as the Covid situation improved, I had the opportunity to obtain two international mobility fellowships from the Université Paris Cité and from the Institut de physique du globe de Paris to carry out an international research project at the University of Perugia in Italy, in the Department of Physics and Geology, hosted by Diego Perugini and Daniele Morgavi. This allowed me to carry out a study that was not initially in my PhD subject. The aim of this study was to investigate the mixing textures and magma mixing timescales between two endmembers, to compare with what occurred prior to the 2010-2013 eruption of Kizimen, based on previous experimental studies [START_REF] Morgavi | The Grizzly Lake complex (Yellowstone Volcano, USA): Mixing between basalt and rhyolite unraveled by microanalysis and X-ray microtomography[END_REF](Morgavi et al., , 2013b(Morgavi et al., , 2013a;;Perugini et al., 2015b). Indeed, magma mixing evidence was shown for this eruption (~50 years before the eruption (Auer et al., 2018)), with a last mingling process prior eruption, leading to the eruption of banded andesites and dacites and andesites. Two original high-temperature magma mixing experiments were performed in Perugia during a 3-week mission in November 2021 and then during a 2-week mission in January 2022.

My thesis therefore focuses on unmonitored ancient Plinian eruptions and on a large pumiceous eruption named "ignimbritic eruption" of a volcano in the island of Dominica, the island with the largest magmatic production in the Lesser Antilles arc, and also on two recent eruptions of two volcanoes in Kamchatka (Bezymianny and well-monitored eruption of Kizimen).

Zoned orthopyroxenes crystals are the main minerals studied throughout the thesis, but I also looked at zoned magnetites, to decipher the timescales of magma residence and magmatic processes history. We also looked at whole-rock compositions and at some melt inclusions and residual glasses to understand better the storage conditions of the magmas. An experimental study is also performed on Kizimen samples, to have a complete view of the system, using natural crystals from volcanic samples and experimental and natural samples in experiments.

This thesis manuscript is divided in 7 chapters. All chapters are introduced in French and the main Chapters are written in the form of peer-reviewed articles, followed by their Supplementary Materials. An appendix (Appendix A) presents a study on two ignimbritic eruptions (33 ka and 51 ka) generated by two volcanoes on the island of Dominica, that was published in November 2020 in Frontiers in Earth Science and that I co-authored [START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF].

Chapter I is the introduction of this thesis, explaining the motivation of our study and provides some context and overview on magmas origin, their evolution, pre-eruptive dynamics and timescales, as well as volcanic eruptions.

In Chapter II, the geological contexts of Dominica and Kamchatka are provided, with a focus on the volcanoes and eruptions studied in my thesis: Morne Trois-Piton Micotrin (Dominica) and the eruptions of Grand Fond (24 ka), Goodwill-Entrance (18 ka), PPR1 (13 ka), PPR2 (13 ka), PPR3 (9 ka). Grand Fond, Goodwill Entrance, PPR1, PPR2 and PPR3, as well as the 1956 climactic eruptive phase of Bezymianny and the 2010-2013 eruption of Kizimen (Kamchatka).

In Chapter III, the sampling methodology and preparation, as well as the analytical techniques are presented. Diffusion chronometry on zoned crystals are presented, along with the state of the art on this topic to understand the timescales of pre-eruptive processes, as well as the strategy of the modelling used here and the uncertainties associated to the method.

The Chapter IV is presented in the form of the article accepted and published in 2021 in Scientific reports [START_REF] Ostorero | Time -window into the transcrustal plumbing system dynamics of Dominica (Lesser Antilles)[END_REF], as well at its Supplementary Materials. Several Plinian eruptions and a larger ignimbritic eruption (Grand Fond) of a central volcano on the island of Dominica were studied.

Two main objectives motivated this study, firstly to determine whether there was a systematic pattern to the estimated pre-eruptive timescales and pre-eruptive dynamics for Plinian eruptions whose magmas are stored at the same depth and secondly, whether the estimated times for larger, deeper stored eruptions are different from those for Plinian eruptions. Combining diffusive timescales and the evolution of magmatic environments, we show that in the 10 to 30 years preceding the four most recent eruptions, two magmatic environments interacted with increasing intensity. The interactions involving the Grand-Fond magmas started about 100 years prior to the eruption. These timescales results can be useful for the monitoring of reactivation signals recorded on the surface, in order to anticipate the likely time remaining before an eruption.

In Chapter V, the 1956 Bezymianny flank collapse is studied, it generated a laterally directed blast and Plinian phase. This paroxysmal phase of the eruption is done through the combined study of zoned orthopyroxenes, magnetites and melt inclusions, as well as residual glasses. The chapter is presented in an article form and is a first version of a manuscript that we aim to submit to Journal of Volcanology and Geothermal Research. We also studied the effect of changing some parameters, that are at stake when modelling diffusion timescales, such as temperature, oxygen fugacity and the diffusion coefficient parametrization. This study shows the importance of studying several minerals from the same eruption.

The Chapter VI is about the 2010-2013 eruption of Kizimen and is subdivided in 4 subchapters.

The first one, Chapter VI.1 is focused on the correlation between diffusive timescales in zoned orthopyroxenes with seismic signals prior to the eruption of Kizimen. We show that the zoned orthopyroxenes record magma mixing processes beginning 1.5 years before the eruption, which is correlated to the pre-eruptive seismic crisis that occurred at the same time. The chapter is presented in a revised peer-reviewed format accepted in October 2022 in Communications, Earth & Environment.

Chapter VI.2 is about the combined study of the zoned opx from Chapter VI.1 and a systematic study of the evolution of magmatic environments where these crystals grew, with zoned magnetites and plagioclase from the same eruption. Timescales of Ti and Mg diffusion were modelled in the zoned magnetites and plagioclase, respectively and compared to the orthopyroxenes, showing shorter timescales. This chapter is presented in an article format and a first version of a manuscript that we aim to submit to Contributions to Mineralogy and Petrology.

Chapter VI. 3 is about the experimental study on Kizimen eruption. We performed for the first-time chemical mixing at sub-liquidus conditions (at which crystallization is promoted) via a chaotic mixing experiment in order to constrain the timescales of the pre-eruptive mixing event prior to the 2010 Kizimen eruption. Crystal Size Distribution methods are applied to the experimental samples that we obtained, as well as in the natural banded andesites of the eruption. We show short timescales associated to magma mixing processes, on the order of minutes to hours. This chapter is written in an article format and is a first version of a manuscript that we aim to submit to Earth and Planetary Science Letters.

Chapter VI.4 is a summary of the second high-temperature magma mixing experiment performed between two natural samples of Kizimen (a basalt and a dacite) (not in the form of a peer-reviewed article).

Finally, Chapter VII is the final chapter and presents a synthesis of the thesis, criticizes the results obtained and provides perspectives for future research.

Résumé

Dans ce deuxième chapitre, le phénomène de subduction puis les contextes géologiques des zones étudiées, les Petites Antilles et l'île de la Dominique ainsi que la péninsule du Kamchatka, à l'est de la Russie sont présentés. Les contextes tectoniques et structuraux des zones d'études sont présentés, ainsi que les compositions des magmas, la structure de la croûte et les volcans d'intérêt. Le chapitre est subdivisé en trois grandes parties.

Premièrement, les zones de subduction lithosphériques proviennent de l'enfoncement d'une lithosphère océanique dans le manteau asthénosphérique sous une autre lithosphère océanique plus jeune ou continentale. Ce phénomène est basé sur la différence de densité entre les plaques lithosphériques, en effet, la lithosphère océanique, s'éloignant de la dorsale, refroidit et devient alors plus dense : elle plonge dans le manteau asthénosphérique [START_REF] Renard | Eléments de géologie -15e édition du Pomerol[END_REF]. Des manifestations magmatiques caractérisent les zones de subduction, générant un volcanisme d'arc (ou orogénique) de type calco-alcalin. Différents modèles permettent d'expliquer la formation de magmas sous les fronts volcaniques et sont détaillés dans le chapitre. Au niveau d'une zone de subduction, les magmas sont de composition basaltique à rhyolitique en passant par les andésites et dacites. Leurs proportions peuvent dépendre de plusieurs facteurs, comme le type de subduction ou le taux de fusion partielle des péridotites [START_REF] Renard | Eléments de géologie -15e édition du Pomerol[END_REF].

En deuxième partie, l'arc des Petites Antilles est abordé. C'est un arc volcanique de 19 îles situé à l'est de la mer des Caraïbes qui provient de la subduction des plaques Nord et Sud-américaines sous la plaque Caraïbe (Macdonald et al., 2000;[START_REF] Wadge | Segmentation of the Lesser Antilles subduction zone[END_REF]. La vitesse de subduction a été estimée à 2 cm/an de 2,4 Ma à aujourd'hui, ce qui correspond à un lent taux de subduction [START_REF] Macdonald | Inversion of magnetic anomalies and sea-floor spreading in the Cayman Trough[END_REF]Macdonald et al., 2000). Un élément important est la subdivision de l'arc au nord de la Martinique. À l'est, d'Anguilla à la Dominique, se trouvent les îles ayant un volcanisme ancien dont l'activité éruptive date de l'Eocène jusqu'à la moitié de l'Oligocène et dont certaines sont maintenant érodées et recouvertes par des plateformes carbonatées. À l'ouest, de Saba à la Dominique, se trouvent les îles volcaniques actives des Caraïbes du Miocène supérieur jusqu'à aujourd'hui (depuis 20 Ma) (Bouysse and Westercamp, 1990;Macdonald et al., 2000). Le volcanisme d'âge inférieur à 2 Ma est localisé sur une zone d'à peu près 10 km de large au niveau de trois segments : le segment nord (de Saba à Montserrat), le segment central (de la Guadeloupe à la Martinique) et le segment sud (Sainte Lucie jusqu'à Grenade) [START_REF] Wadge | Segmentation of the Lesser Antilles subduction zone[END_REF]. 16 volcans de cet arc ont connu une activité durant l'Holocène (depuis 12 000 ans) et 9 d'entre eux sont entrés en éruption durant la période historique (depuis 1270) [START_REF] Boudon | Volcanological evolution of Montagne Pelée (Martinique): A textbook case of alternating Plinian and dome-forming eruptions[END_REF][START_REF] Brown | Report IV of the GVM/IAVCEI contribution to the Global Assessment Report on Disaster Risk Reduction[END_REF][START_REF] Leone | Integrating spatial accessibility in the design of volcano evacuation plans in the French West Indies (Guadeloupe and Martinique)[END_REF]. Ils émettent des magmas différenciés avec un haut potentiel explosif. La nature de la croûte sous les volcans de l'arc des Petites Antilles est peu connue, même si les roches de la croûte interagissent avec les magmas lors de leur ascension et stockage [START_REF] Boudon | Volcanological evolution of Montagne Pelée (Martinique): A textbook case of alternating Plinian and dome-forming eruptions[END_REF]. Des études géophysiques ont permis d'identifier un système crustal entre 22 et 37 km de profondeur composé de trois couches [START_REF] Arnaiz-Rodríguez | La estructura cortical del arco de las Antillas Menores estimada a partir de la técnica de funciones receptoras[END_REF][START_REF] Boynton | A seismic refraction investigation of crustal structure beneath the Lesser Antilles island arc[END_REF][START_REF] Kopp | Deep structure of the central Lesser Antilles Island Arc: Relevance for the formation of continental crust[END_REF][START_REF] Melekhova | Lateral variation in crustal structure along the Lesser Antilles arc from petrology of crustal xenoliths and seismic receiver functions[END_REF]. La structure de la croûte est cohérente avec de la différenciation magmatique ayant lieu au sein de réservoirs verticaux dominés par des zones de mush magmatique, où des lentilles de magmas sont stockées à différentes profondeurs, avec des liquides interstitiels piégés dans des parties cristallisées (Cashman et al., 2017;[START_REF] Melekhova | Lateral variation in crustal structure along the Lesser Antilles arc from petrology of crustal xenoliths and seismic receiver functions[END_REF][START_REF] Melekestsev | Experimental Evidence for Polybaric Differentiation of Primitive Arc Basalt beneath St. Vincent, Lesser Antilles[END_REF][START_REF] Parman | Along-Arc Variations in the Pre-Eruptive H2O Contents of Mariana Arc Magmas Inferred from Fractionation Paths[END_REF]. Il existerait, sous pratiquement toutes les îles de l'arc, une zone de stockage transcrustale entre 5 et 9 km de profondeur avec un système de mush étendu verticalement et souvent latéralement (Balcone-Boissard et al., 2018;[START_REF] Christopher | Crustal-scale degassing due to magma system destabilization and magma-gas decoupling at Soufrière Hills Volcano, Montserrat. Geochemistry[END_REF][START_REF]Evidence for an Active, Transcrustal Magma System in the Last 60 ka and Eruptive Degassing Budget (H 2 O, CO 2[END_REF][START_REF] Edmonds | Extensive, water-rich magma reservoir beneath southern Montserrat[END_REF][START_REF] Martel | Magma storage conditions and control of eruption regime in silicic volcanoes: Experimental evidence from Mt[END_REF][START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF][START_REF] Pichavant | Experimental Parametrization of Magma Mixing: Application to the ad 1530 Eruption of La Soufrière, Guadeloupe (Lesser Antilles)[END_REF][START_REF] Pichavant | Physical conditions, structure, and dynamics of a zoned magma chamber: Mount Pelée (Martinique, Lesser Antilles Arc)[END_REF]Solaro et al., 2019).

L'île de la Dominique appartient à la partie décente du nord de l'arc. Cette île a la particularité d'avoir le plus grand nombre de volcans actifs de l'arc (quatre volcans, comparé à un volcan actif en général pour les autres îles de l'arc) (Lindsay et al., 2005c). Le volcan de cette île étudié ici est le Morne Trois-Piton Micotrin. Ce volcan a connu trois éruptions ponceuses de grand volume dites « ignimbritiques » (Grand Bay > 60 ka, Roseau 33 ka et Grand Fond 24 ka) (Boudon et al., 2017;Newhall and Self, 1982), qui ont émis chacune un volume de magma de l'ordre de 5 km 3 « Dense Rock Equivalent » (DRE). Ces éruptions débutent par une phase Plinienne générant une colonne éruptive convective produisant des retombées de ponces et cendres et se poursuivant par la mise en place d'importants volumes d'écoulements pyroclastiques. Puis, une série de 8 éruptions Pliniennes, mettant en jeu des volumes de magma plus faibles d'au moins un ordre de grandeur (0,5 km 3 ), se sont succédées principalement après la dernière éruption ignimbritique de Grand Fond (Boudon et al., 2017;Newhall and Self, 1982). Le travail réalisé dans cette thèse a porté sur cinq éruptions du volcan Morne Trois Piton-Micotrin, l'éruption ignimbritique de Grand Fond et 4 des 8 éruptions Pliniennes postérieures à Grand Fond : Goodwill Entrance : 18 095 ± 205 ans cal BP, « Plinienne Post Roseau » 1 (PPR1) : 12 619 ± 88 ans cal BP, PPR2 : 12 513 ± 129 ans cal BP et PPR3 : 9 391 ± 94 ans cal BP. Concernant les profondeurs de stockage des magmas, les magmas ayant généré des éruptions ignimbritiques se situent à la limite entre la croûte inférieure et la croûte moyenne dans les modèles de la structure de la croûte de l'arc des Petites Antilles [START_REF] Kopp | Deep structure of the central Lesser Antilles Island Arc: Relevance for the formation of continental crust[END_REF][START_REF] Melekhova | Lateral variation in crustal structure along the Lesser Antilles arc from petrology of crustal xenoliths and seismic receiver functions[END_REF]. Pour les magmas se trouvant à des profondeurs moins importantes (magmas ayant généré les éruptions Pliniennes), ces magmas se situeraient à la limite entre la croûte moyenne et la croûte supérieure.

En dernière partie, le contexte géologique et structural du Kamchatka est présenté. La région du Kamchatka fait partie de la ceinture de feu du Pacifique, qui est une longue chaîne de 40 000 km de long avec plus de 450 volcans actifs produisant des éruptions explosives et effusives, ainsi que des volcans dormants [START_REF] Eichelberger | Preface, in: Volcanism and Subduction: The Kamchatka Region[END_REF]. La péninsule du Kamchatka se situe au-dessus de la marge nord-ouest de la plaque Pacifique plongeant à une vitesse de ~8 cm/an [START_REF] Demets | Oblique convergence and deformation along the Kuril and Japan Trenches[END_REF], formant l'arc volcanique du Kamchatka. L'âge de la subduction date d'à peu près 95 Ma à la pointe sud du Kamchatka (Portnyagin et al., 2007a;[START_REF] Renkin | Depth and age in the north Pacific[END_REF]. La structure moderne du Kamchatka ne s'est formée que durant les derniers millions d'années [START_REF] Lander | The origin of the modern Kamchatka Subduction Zone[END_REF]. Le volcanisme datant du Pléistocène supérieur-Holocène au Kamchatka forme deux chaînes : la Eastern volcanic belt (ceinture volcanique de l'est) et la Sredinny Range. La ceinture de l'est peut être subdivisée en deux : le front volcanique de l'est (Eastern Volcanic Front ou EVF) et la zone volcanique de Dépression Centrale du Kamchatka (Central Kamchatka Depression ou CKD) [START_REF] Ponomareva | Late Pleistocene-Holocene volcanism on the Kamchatka Peninsula, Northwest Pacific Region, in: Volcanism and Subduction: The Kamchatka Region[END_REF]. L'activité récente se concentre au niveau des centres volcaniques vigoureux du Klyuchevskoy et du Shiveluch dans la CKD, où la plus grosse production magmatique est enregistrée. Les principales sources contribuant à la diversité des magmas de la CKD sont : le coin du manteau, la plaque en subduction et possiblement une plaque lithosphérique chevauchante (Portnyagin et al., 2007b). Les changements géochimiques le long du segment nord du Kamchatka (en éléments majeurs et traces, radiogéniques et ratios d'isotopes stables dans les laves primitives) peuvent être expliqués par le déchirement de la plaque Pacifique en subduction sous la CKD, divisant la CKD en deux parties où les conditions de génération de magma changent, allant d'une composition typique d'arc de subduction au sud à une composition de volcanisme intraplaque ou océanique au nord. Les magmas parents des volcans du sud auraient 3 principales sources : manteau déplété, liquides éclogitiques et des fluides du slab ou composantes lithosphériques (Portnyagin et al., 2007a). Des études tomographiques et sismiques ont montré que certains volcans de la CKD (dont le Bezymianny) partagent une source commune de fluides magmatiques provenant d'une chambre magmatique située à 30 km de profondeur [START_REF] Koulakov | Three different types of plumbing system beneath the neighboring active volcanoes of Tolbachik, Bezymianny, and Klyuchevskoy in Kamchatka[END_REF][START_REF] Shapiro | Deep and shallow long-period volcanic seismicity linked by fluid-pressure transfer[END_REF] Enfin, les volcans du Bezymianny et Kizimen, principaux volcans d'étude de cette thèse, sont présentés.

Le Bezymianny fait partie de la CKD. C'est un stratovolcan andésitique datant de l'Holocène (11 000 ans) [START_REF] Ponomareva | Late Pleistocene-Holocene volcanism on the Kamchatka Peninsula, Northwest Pacific Region, in: Volcanism and Subduction: The Kamchatka Region[END_REF]. Il est devenu célèbre lors de sa première éruption historique et de sa phase paroxysmale du 30 mars 1956, après une période de dormance d'un millénaire (Braitseva et al., 1991). La phase explosive du 30 mars 1956 a été précédée par six mois d'activité volcanique dominés par des explosions Vulcaniennes, la croissance d'un dôme de lave à son sommet puis la croissance d'une cryptodôme dans le flanc est de l'édifice. Un glissement du flanc du volcan à grande échelle (avalanche de débris de 0,5 km 3 sur 22 km) a entraîné la dépressurisation rapide du cryptodôme et une forte explosion le 30 mars 1956 [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF]. La puissante explosion dirigée latéralement a détruit la partie orientale de l'édifice volcanique sur 500 km 2 , produisant entre 0,2 et 0,4 km 3 de dépôts, et a été suivie par une phase Plinienne générant des retombées de cendres et de ponces ainsi que des courants de densité pyroclastiques ponceux d'un volume de 0,2-0,3 km 3 à 0,5 km 3 , respectivement, couvrant 40 km 2 [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF]. Au cours des 75 dernières années, le volcan Bezymianny a montré une diversité de styles d'éruptions, avec une dualité entre les éruptions Vulcaniennes et les éruptions à croissance de dômes de lave générant des courants de densité pyroclastiques de cendres et de blocs. Concernant l'architecture du système magmatique du Bezymianny, il serait constitué de plusieurs niveaux et contiendrait au moins deux réservoirs magmatiques dans la croûte (200-300 MPa à 900 ± 5 0°C et 500-750 MPa à 975 ± 25 °C, correspondant à des profondeurs de l'ordre de 7 et 18 km) (études pétrologiques, tomographiques et sismiques) [START_REF] Davydova | Petrology of mafic enclaves in the 2006-2012 eruptive products of Bezymianny Volcano, Kamchatka[END_REF]Shcherbakov and Plechov, 2010 ;Shcherbakov et al., 2011 ;[START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]. Des épisodes de recharge du réservoir le moins profond par du magma saturé en éléments volatils est un des mécanismes possiblement déclencheur d'une éruption, maintenant une température stable dans la chambre, entre 940 et 980°C [START_REF] Davydova | Petrology of mafic enclaves in the 2006-2012 eruptive products of Bezymianny Volcano, Kamchatka[END_REF]. D'après le modèle de Turner et al. (2013), expliquant les dynamiques pré-éruptives avant la phase paroxysmale de l'éruption de 1956, des mélanges de magmas ont eu lieu, avec une injection de magma basique dans le premier réservoir, suivi d'une interaction entre cette poche de magma et le réservoir plus superficiel.

Le Kizimen est un volcan datant du Pléistocène Supérieur-Holocène inférieur faisant aussi partie de la CKD. Quatre cycles d'activités ont été identifiés au Kizimen, datés de 12-11 ka à aujourd'hui, avant l'éruption récente de 2010, grâce à une étude téphrochronologique (Melekestsev et al., 1995). Chaque cycle est caractérisé par des périodes de forte activité suivis par une phase de déclin puis par une phase de repos. Chaque cycle commence par de larges explosions et se termine par la formation d'un dôme de lave et l'émission de coulées de lave visqueuses et épaisses (Melekestsev et al., 1995).

En 1963, le volcan a été le siège d'une intense activité sismique (séismes jusqu'à 35 km de profondeur ; [START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF]). En 2010-2013, il a produit sa première éruption magmatique historique avec l'émission de 0,4 km 3 Dense Rock Equivalent (DRE) de magma. L'éruption a débuté le 11 novembre 2010 et a été précédée par un an et demi de crise sismique [START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF]. Une phase explosive générant des courants de densité pyroclastique a été suivie par l'extrusion d'un dôme de lave qui a progressivement formé une coulée de lave très épaisse (Auer et al., 2018). De nombreux courants de densité pyroclastiques de blocs et de cendres ont été produits suite à l'effondrement gravitaire du dôme de lave et du front de la coulée. Un modèle proposé par Auer et al. (2018) pour l'éruption de 2010 a suggéré qu'une injection de magma basaltique dans un réservoir dacitique s'était produite en profondeur en 1963, associée à la crise sismique. Puis, après une longue période de dormance, une phase d'activité sismique a débuté en avril 2009 et s'est poursuivie pendant une année et demie avant l'éruption [START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF]. Un mélange rapide des composants du magma dans un réservoir probablement stratifié serait à l'origine de l'éruption en novembre 2010, éjectant un magma mal mélangé attesté par l'abondance d'andésites et de dacites rubanées. Concernant les profondeurs de stockage des magmas, ils seraient stockés entre 5 et 11 km (~810 °C et ~123 MPa (~5 km)) (Auer et al., 2018).

La surveillance de ces deux volcans est opérée au Kamchatka par le département de l'étude géophysique de l'Académie des Sciences de Russie, à travers un réseau de stations sismiques permanentes mais aussi par l'étude des déformations par stations GPS et des études géodésiques (InSAR, images de webcam...).

In this second chapter, the geological contexts of the studied volcanoes, in the subduction zones of the Lesser Antilles (Dominica island) and the Kamchatka peninsula in eastern Russia are presented.

II.1. Subduction zone and associated volcanism

The phenomenon of lithospheric subduction is the sinking of an oceanic lithosphere under another younger oceanic or continental lithosphere (Renard et al., 2015) (Figure II.1a). This process is based on the difference in density between the lithospheric plates: as the oceanic lithosphere moves away from the ridge and cools down, it becomes denser and can plunge into the asthenospheric mantle. These processes are involved in the recycling of the oceanic lithosphere, with the subduction of part of the oceanic sediments as well as the recycling of chemical elements such as volatiles through thermal, geochemical and petrological exchange processes [START_REF] Kerrick | Metamorphic devolatilization of subducted oceanic metabasalts: implications for seismicity, arc magmatism and volatile recycling[END_REF][START_REF] Rüpke | Serpentine and the subduction zone water cycle[END_REF][START_REF] Zellmer | Volatiles in subduction zone magmatism[END_REF]. Different types of subduction exist: oceanic or continental subductions, oceancontinent subductions (as in Kamchatka, where the oceanic Pacific plate goes under Kamchatka's peninsula) or ocean-ocean subductions (as in the Lesser Antilles arc where the North and South Atlantic plates go under the Caribbean plate) [START_REF] Renard | Eléments de géologie -15e édition du Pomerol[END_REF]. Subduction between two continents (continent-continent, as in Himalaya) or between a continent and an ocean ("obduction", as in Oman) are also evidenced [START_REF] Renard | Eléments de géologie -15e édition du Pomerol[END_REF].

Subduction zones generate voluminous magma and arc (or orogenic) volcanism of calc-alkaline type and can be organized as arcs of islands or chains of mountains [START_REF] Manning | The chemistry of subduction-zone fluids[END_REF] (Figure II.1a). Only 3 to 4% of the magma is erupted while the majority is stored in the crust [START_REF] Paterson | Arc Magmatic Tempos: Gathering the Evidence[END_REF]. Depending on the type of subduction, the magmas are more or less differentiated, with the production of small volumes of andesites to large volumes of rhyolites where crustal processes are dominant, through silicic caldera-forming eruptions (~1 km 3 Dense Rock Equivalent (DRE) magma volumes for pre-historic basaltic andesites-andesitic eruptions of Montagne Pelée (Martinique) compared to ~10 5 km 3 for the large silicic volcanic province in Sierra Madre Occidental (Mexico), for example) [START_REF] Bryan | New Insights into Crustal Contributions to Large-volume Rhyolite Generation in the Mid-Tertiary Sierra Madre Occidental Province, Mexico, Revealed by U-Pb Geochronology[END_REF][START_REF] Carazzo | Transition from stable column to partial collapse during the 79 cal CE P3 Plinian eruption of Mt. Pelée volcano[END_REF][START_REF] Carazzo | Marginally stable recent Plinian eruptions of Mt. Pelée volcano (Lesser Antilles): the P2 AD 280 eruption[END_REF]. In island arcs, magmas are less differentiated (basalts, basic andesites) and andesites are mainly found, as in Dominica or Martinique (Lindsay et al., 2005b;[START_REF] Smith | Mt. Pelée, Martinique : a study of an active island-arc volcano[END_REF], and there is little involvement of the continental crust.

Different models can explain the formation of magmas under volcanic fronts. These magmas are generated above the plunging plate generally at 110 km depth, except for arcs with a high subduction angle (> 70°, in Vanuatu and the Solomon Islands) or a low subduction velocity (< 1 cm/year at the Aeolian arc) [START_REF] Tatsumi | Migration of fluid phases and genesis of basalt magmas in subduction zones[END_REF]. The geothermal gradient of a current subduction zone leads to the destabilization of hydrated phases of the plunging oceanic lithosphere (previously hydrated during hydrothermal processes at the ridges), at depths below 100 km, such as serpentines, amphiboles, chlorites, as well as oceanic sediments, as a result of pressure and temperature changes (Figure II.1a). Following their dehydration, fluids concentrate in a zone of the overlying mantle called the "mantle wedge", around 100 km depth, at the lithosphere-asthenosphere boundary below the arc. The plunging panel enriched in incompatible elements reacts with the overlying mantle materials, hydrating the peridotite minerals [START_REF] Tatsumi | Formation of the volcanic front in subduction zones[END_REF]. With convection in the mantle wedge, these peridotites are transported at depth on the slab, resulting in destabilization of the amphiboles around 110 km (30 kbar) [START_REF] Schmidt | Experimentally based water budgets for dehydrating slabs and consequences for arc magma generation[END_REF][START_REF] Tatsumi | Formation of the volcanic front in subduction zones[END_REF], just below the volcanic front. These fluids are classically mainly composed of H2O, with Na and Cl, low Si content and a modest load of alkali and alkaline earth metals that modify the mineralogy of the mantle, by metasomatosis [START_REF] Manning | The chemistry of subduction-zone fluids[END_REF]. The mantle can therefore enter into partial melting (1000-1100 °C) [START_REF] Tatsumi | Formation of the volcanic front in subduction zones[END_REF] when the hydrated peridotite solidus intersects the geotherm of the mantle corner, around 100 km with melt rates on the order of 15-20% [START_REF] Jarrard | Subduction fluxes of water, carbon dioxide, chlorine, and potassium[END_REF][START_REF] Manning | The chemistry of subduction-zone fluids[END_REF][START_REF] Peacock | Fluid Processes in Subduction Zones[END_REF][START_REF] Tatsumi | Migration of fluid phases and genesis of basalt magmas in subduction zones[END_REF][START_REF] Tatsumi | Subduction Zone Magmatism[END_REF]. Hydrated primary magmas rise to the surface by density difference with the surrounding unmelted peridotite. These magmas interact with the mantle and then with the crust, leading to magma mixing and/or crustal contamination in the case of storage at the crust/mantle interface. These magmas can also differentiate during a more or less long storage in magmatic reservoirs.

However, the discovery of rocks called adakites in subduction zones such as in the Aleutians (Adak Island, Alaska) [START_REF] Kay | Aleutian magnesian andesites: Melts from subducted Pacific ocean crust[END_REF] and north of Kamchatka (Shiveluch) (Volynets et al., 1999b) calls into question this first model for the origin of arc magmas. Their compositions are close to dacites and characterized by a high-water content, an enrichment in Si and a low concentration of K and heavy rare earths elements compared to basalts and basalt-andesites of classical subduction zones. One model to interpret the formation of these magmas is that they are derived from partial melting of the plunging plate (and not the overlying mantle) [START_REF] Renard | Eléments de géologie -15e édition du Pomerol[END_REF]. These magmas are present in 34 active island arcs and in almost all those where the plunging oceanic crust is less than 40-50 Ma old [START_REF] Renard | Eléments de géologie -15e édition du Pomerol[END_REF]. Nevertheless, the former model remains valid in the context of relatively old oceanic lithosphere subductions.

Generally, the primary arc magmas are hydrated and rich in incompatible elements (Ba, Rb, Th, K), which explains why they do not have a basaltic composition but can evolve between basaltic-andesitic compositions (Figure II.1b). Their geochemical compositions differ from mid-ocean ridge basalts (MORB) and Ocean Island basalts (OIB) generated in other geological contexts (Macdonald et al., 2000) (Figure II.1b). They show a negative anomaly in niobium (Nb) and tantalum (Ta). The metasomatism of the mantle wedge is found in the composition of primary magmas, with radiogenic isotopic ratios intermediate between those of MORB and oceanic sediments, such as 87 Sr/ 86 Sr.

H2O abundance in arc magmas was previously estimated to be around 1.3-7.2 wt% in melt inclusions hosted in olivines, however, as diffusion of hydrogen in melt inclusions is fast, the initial H2O content of arc magmas in the lower crust may not be accurately recorded [START_REF] Annen | The Genesis of Intermediate and Silicic Magmas in Deep Crustal Hot Zones[END_REF][START_REF] Bucholz | Post-entrapment modification of volatiles and oxygen fugacity in olivine-hosted melt inclusions[END_REF][START_REF] Gaetani | Rapid reequilibration of H2O and oxygen fugacity in olivine-hosted melt inclusions[END_REF][START_REF] Plank | Why do mafic arc magmas contain ∼4wt% water on average? Earth Planet[END_REF]Portnyagin et al., 2008;[START_REF] Urann | High water content of arc magmas recorded in cumulates from subduction zone lower crust[END_REF]. Recent measurements on lower crustal cumulates, that crystallized in the crust at a pressure high enough to record the original water content, show higher water contents of 12-20 wt% H2O, as predicted by phase equilibria studies but absent from previous melt inclusions studies [START_REF] Goltz | Evidence for superhydrous primitive arc magmas from mafic enclaves at Shiveluch volcano, Kamchatka[END_REF][START_REF] Grove | Magnesian andesite and dacite lavas from Mt. Shasta, northern California: products of fractional crystallization of H2O-rich mantle melts[END_REF][START_REF] Krawczynski | Amphibole stability in primitive arc magmas: effects of temperature, H2O content, and oxygen fugacity[END_REF][START_REF] Urann | High water content of arc magmas recorded in cumulates from subduction zone lower crust[END_REF]. Copyrighted material, unauthorised distribution (Figure 1; [START_REF] Stern | SUBDUCTION ZONES[END_REF]

II.2. The Lesser Antilles arc

II.2.1. Tectonic and structural context

The Lesser Antilles arc is a volcanic arc of 19 islands located in the eastern Caribbean Sea. It extends from the island of Saba in the north to the island of Grenada in the south (800 km long), from 12° to 81°N, with a convexity towards the east (Macdonald et al., 2000;[START_REF] Wadge | Segmentation of the Lesser Antilles subduction zone[END_REF]. It has been active since 40 Ma and was formed as a result of the curved and oblique subduction of the northern and southern American plates beneath the Caribbean plate (Bouysse and Westercamp, 1990;[START_REF] Feuillet | Tectonic context of moderate to large historical earthquakes in the Lesser Antilles and mechanical coupling with volcanoes[END_REF]Macdonald et al., 2000;[START_REF] Symithe | Current block motions and strain accumulation on active faults in the Caribbean[END_REF] (Figure II.2a). An important feature is the subdivision of the arc north of Martinique. To the east, from Anguilla to Dominica, there are islands with ancient volcanism whose eruptive activity dates from the Eocene to mid-Oligocene . Some of these islands are now eroded and covered by carbonate platforms. To the west, from Saba to Dominica, are the active volcanic islands from the Recent Miocene to the present (since 20 Ma) (Bouysse and Westercamp, 1990;Macdonald et al., 2000) (Figure II.2b). This westward shift in activity is debated but is surely due to a decrease in the dip of the plunging plate due to subduction of parts of an aseismic ridge (the Barracuda Ridge) [START_REF] Bouysse | The Lesser Antilles Island Arc: Structure and Geodynamic Evolution, in: Initial Reports of the Deep Sea Drilling Project[END_REF]Macdonald et al., 2000;[START_REF] Mccann | Subduction of aseismic ridges beneath the Caribbean Plate: Implications for the tectonics and seismic potential of the northeastern Caribbean[END_REF]. This event would have halted the subduction process, causing the arc to migrate to the forearc basin of the subduction zone [START_REF] Boudon | Volcanological evolution of Montagne Pelée (Martinique): A textbook case of alternating Plinian and dome-forming eruptions[END_REF][START_REF] Bouysse | The Lesser Antilles Island Arc: Structure and Geodynamic Evolution, in: Initial Reports of the Deep Sea Drilling Project[END_REF]Macdonald et al., 2000;[START_REF] Mccann | Subduction of aseismic ridges beneath the Caribbean Plate: Implications for the tectonics and seismic potential of the northeastern Caribbean[END_REF]. Since the Upper Oligocene-Late Miocene, in the south of Martinique, the arc has remained in a stable position and volcanic activity has been continuous, with a slight westward migration (less than 10 km, the width of the island) while in the north of this island, volcanic activity stopped during this period and migrated westward to its present position (Macdonald et al., 2000;[START_REF] Sigurdsson | Marine Tephrochronology and Quaternary Explosive Volcanism in the Lesser Antilles Arc[END_REF][START_REF] Wadge | Segmentation of the Lesser Antilles subduction zone[END_REF]. A revised tectonic model for the eastern Caribbean based on a compilation of marine magnetic anomaly and seismic data suggests an arc migration in both directions [START_REF] Allen | The role of arc migration in the development of the Lesser Antilles: A new tectonic model for the Cenozoic evolution of the eastern Caribbean[END_REF] (Figure II.3). Indeed, [START_REF] Allen | The role of arc migration in the development of the Lesser Antilles: A new tectonic model for the Cenozoic evolution of the eastern Caribbean[END_REF] propose that two phases of back-arc spreading and eastward-arc migration toward the incoming Atlantic, and a third phase of arc migration to the west, subdivided the earlier back-arc basin on both sides of the Lesser Antilles arc (Figure II.3). The two branches come together at Dominica to form a single island chain in the south where the volcanic deposits of the old arc are frequently overlain by those of the recent arc. This part of the arc is bordered to the west by the Grenada back-arc basin (3 km deep).

The rate of subduction has been estimated at 2 cm/yr from 2.4 Ma to present and 4 cm/yr between 8.3 and 2.4 Ma (Macdonald and Holcombe, 1978;Macdonald et al., 2000), which corresponds to a slow rate of subduction compared to other arcs, for example the Java arc (8.1 cm/yr) [START_REF] Jarrard | Relations among subduction parameters[END_REF]. These low velocities are thought to account for the low magmatic production of the arc, which were, for the last 300, 10,000, and 100,000 years, of 5, 3, and 4 km 3 Ma -1 .km -1 , respectively (Macdonald et al., 2000;Wadge, 1984). From seismic data, [START_REF] Wadge | Segmentation of the Lesser Antilles subduction zone[END_REF] established the configuration of the Wadati-Benioff zone beneath the Lesser Antilles. The arc has a direction of 330° and a dip of 50-60° for the arc segment north of Martinique, whereas south of St. Lucia, the direction is 20° and the dip of the dipping panel is 40-50° and then nearly vertical beneath the island of Grenada (Macdonald et al., 2000;[START_REF] Wadge | Segmentation of the Lesser Antilles subduction zone[END_REF]. This segmentation may represent the deformation of a single subducting plate or two separate plates corresponding to the North and South American plates with a triple point at the Martinique-St. Lucia islands [START_REF] Wadge | Segmentation of the Lesser Antilles subduction zone[END_REF].

Volcanism younger than 2 Ma has focused on an area about 10 km wide in three segments: the northern arc (Saba to Montserrat), the central arc (Guadeloupe to Martinique), and the southern arc (St. Lucia to Grenada) [START_REF] Wadge | Segmentation of the Lesser Antilles subduction zone[END_REF] 16 volcanoes in this arc experienced activity during the Holocene, including 12 on the major volcanic islands of the recent arc and 9 during the historical period (since 1270) [START_REF] Boudon | Volcanological evolution of Montagne Pelée (Martinique): A textbook case of alternating Plinian and dome-forming eruptions[END_REF][START_REF] Brown | Report IV of the GVM/IAVCEI contribution to the Global Assessment Report on Disaster Risk Reduction[END_REF][START_REF] Leone | Integrating spatial accessibility in the design of volcano evacuation plans in the French West Indies (Guadeloupe and Martinique)[END_REF]. These volcanoes are from north to south: Mount Scenery (Saba), The Quill (St. Eustatius), Mount Misery (St. Kitts), Nevis Peak (Nevis), Soufrière Hills (Montserrat), La Soufrière (Guadeloupe), Morne Aux Diables, Morne Diablotins, Morne Trois Pitons-Micotrin and the volcanic complex of Morne Plat Pays (Dominica), Montagne Pelée (Martinique), Soufrière (Sainte Lucia), Soufrière (St. Vincent) and Kick'em Jenny and Kick'em Jack (Grenada) [START_REF] Boudon | Volcanological evolution of Montagne Pelée (Martinique): A textbook case of alternating Plinian and dome-forming eruptions[END_REF].

(Figure II.2b).
Monitoring of the volcanoes of the islands of St. Kitts and Nevis, Dominica, St. Lucia, St. Vincent and the Grenadines and Grenada are operated by the Seismic Research Unit of the University of the West Indies, while the Institut de physique du globe de Paris monitors the volcanic activity in Martinique and Guadeloupe and the Montserrat Volcano Observatory operates in Montserrat (Lindsay et al., 2005a). These agencies have to work closely with civil authorities in case of volcanic unrest. 10 seismic stations of the University of the West Indies are monitoring the volcanoes of Dominica (Lindsay et al., 2005a).

II.2.2. Compositions of the magmas

The rate of magma production varies along the Lesser Antilles arc, with the central arc being the most active and having the largest islands in the arc, including the island of Dominica which has the largest magma production of the entire Caribbean arc: on the order of 40 km 3 for the last 100 ka compared to 5-8 km 3 for the neighboring islands of Martinique and Guadeloupe (with a mean magma production rate of 0.6-0.7 km 3 /1,000 years for Montagne Pelée in Martinique [START_REF] Boudon | Volcanological evolution of Montagne Pelée (Martinique): A textbook case of alternating Plinian and dome-forming eruptions[END_REF]) (Macdonald et al., 2000;Wadge, 1984). In contrast to the central arc, the northern and southern arcs produced smaller volumes of magma (~0-5 km 3 /1000 years) (Macdonald et al., 2000;[START_REF] Metcalfe | Insights into the Origin and Dynamics of Past Eruptions of La Soufrière de Guadeloupe from Melt Inclusions: Implications for Future Unrest and Environmental Impact[END_REF]Wadge, 1984).

Regarding magma compositions, geochemical studies show a range of magma types: the northern arc from Saba to Montserrat, has magmas of andesitic (K-poor tholeiites) to dacitic compositions (low volumes of basalts (~10% on St. Kitts and ~3% on Montserrat and rare rhyolites for St. Kitts and St. Eustatius), the central arc (Guadeloupe to St. Lucia) has magmas that are predominantly andesitic in composition (with a few basalts, dacites, and rare rhyolites) [START_REF] Brown | Geochemistry of the Lesser Antilles volcanic island arc[END_REF][START_REF] Davidson | Isotopic and trace element constraints on the petrogenesis of subduction-related lavas from Martinique, Lesser Antilles[END_REF][START_REF] Hawkesworth | Magma genesis in the lesser Antilles island arc[END_REF]Lindsay et al., 2005c;Macdonald et al., 2000;[START_REF] Rea | The Lesser Antilles[END_REF][START_REF] White | Sediment subduction and magma genesis in the Lesser Antilles; isotopic and trace element constraints[END_REF]. These rocks belong to the calc-alkaline (medium-K) series (Lindsay et al., 2005c). The latter, from St. Vincent to Grenada, consists of basalts and basaltic andesites (in St. Vincent) with some andesites from magma mixtures (Macdonald et al., 2000;[START_REF] Wadge | Segmentation of the Lesser Antilles subduction zone[END_REF].

II.2.3. Structure of the crust

The nature of the crust of volcanoes in the Lesser Antilles arc is poorly understood, although it is known that crustal rocks interact with magmas during their ascent and storage [START_REF] Boudon | Volcanological evolution of Montagne Pelée (Martinique): A textbook case of alternating Plinian and dome-forming eruptions[END_REF]. Geophysical studies have identified a crustal system between 22 and 37 km depth composed of three layers (with the identification of two seismic refractors) [START_REF] Arnaiz-Rodríguez | La estructura cortical del arco de las Antillas Menores estimada a partir de la técnica de funciones receptoras[END_REF][START_REF] Boynton | A seismic refraction investigation of crustal structure beneath the Lesser Antilles island arc[END_REF][START_REF] Kopp | Deep structure of the central Lesser Antilles Island Arc: Relevance for the formation of continental crust[END_REF]Melekhova et al., 2019) (Figure II.3a). The upper layer is composed of volcaniclastic and sedimentary rocks with abundant fractures (P-wave velocity vP < 6 km s -1 ) [START_REF] Kiddle | Crustal structure beneath Montserrat, Lesser Antilles, constrained by xenoliths, seismic velocity structure and petrology[END_REF][START_REF] Kopp | Deep structure of the central Lesser Antilles Island Arc: Relevance for the formation of continental crust[END_REF][START_REF] Melekhova | Lateral variation in crustal structure along the Lesser Antilles arc from petrology of crustal xenoliths and seismic receiver functions[END_REF]. The middle layer of the crust is felsic in nature and the lower layer is composed of plutonic magmatic rocks (with cumulates; vP = 6.9 km s -1 ) [START_REF] Boynton | A seismic refraction investigation of crustal structure beneath the Lesser Antilles island arc[END_REF][START_REF] Melekhova | Lateral variation in crustal structure along the Lesser Antilles arc from petrology of crustal xenoliths and seismic receiver functions[END_REF]. The depth of the Moho has not been directly seismically observed in previous studies but its depth is estimated to be between 24 and 35 km (Arnaiz-Rodríguez et al., 2016;[START_REF] Boynton | A seismic refraction investigation of crustal structure beneath the Lesser Antilles island arc[END_REF][START_REF] Kopp | Deep structure of the central Lesser Antilles Island Arc: Relevance for the formation of continental crust[END_REF][START_REF] Melekhova | Lateral variation in crustal structure along the Lesser Antilles arc from petrology of crustal xenoliths and seismic receiver functions[END_REF]. The crustal structure obtained by [START_REF] Melekhova | Lateral variation in crustal structure along the Lesser Antilles arc from petrology of crustal xenoliths and seismic receiver functions[END_REF] shows variable depths of the middle crust and Moho boundary over short distances (tens of kilometers): between 11 and 32 km and between 25 and 37 km, respectively (Moho at a depth of 25 km beneath St. Eustatius and 37 km beneath St. Kitts and Dominica). The islands of Dominica and Martinique are the islands with the greatest crustal thicknesses (Figure II.3a). These variations in middle layer boundary depths can be explained by variability in estimated magmatic water contents and differentiation depths, also influencing xenolith lithologies and P-wave velocity variations may also reflect changes in magma formation rates and processes in the crust and mantle [START_REF] Melekhova | Lateral variation in crustal structure along the Lesser Antilles arc from petrology of crustal xenoliths and seismic receiver functions[END_REF]. The variation along the arc in crustal structure could reflect upwelling of heterogeneous magmas from the mantle due to variations in water-rich flows from the plunging plate [START_REF] Melekhova | Lateral variation in crustal structure along the Lesser Antilles arc from petrology of crustal xenoliths and seismic receiver functions[END_REF]. The crustal structure is consistent with magmatic differentiation taking place within vertical reservoirs dominated by magmatic mush zones, with little interstitial trapped melt (Cashman et al., 2017;[START_REF] Melekhova | Lateral variation in crustal structure along the Lesser Antilles arc from petrology of crustal xenoliths and seismic receiver functions[END_REF][START_REF] Melekestsev | Experimental Evidence for Polybaric Differentiation of Primitive Arc Basalt beneath St. Vincent, Lesser Antilles[END_REF][START_REF] Parman | Along-Arc Variations in the Pre-Eruptive H2O Contents of Mariana Arc Magmas Inferred from Fractionation Paths[END_REF].

Regarding the storage conditions of magmas in the crust, experimental studies and studies on melt inclusions trapped in crystals allow an estimate of pressures, temperatures and storage depths of magmas. Magmatic storage systems beneath the various volcanic islands of the Lesser Antilles suggests the existence beneath each island of a transcrustal storage zone between 5 and 9 km depth with a vertically and often laterally extensive mush system (Balcone-Boissard et al., 2018;[START_REF] Christopher | Crustal-scale degassing due to magma system destabilization and magma-gas decoupling at Soufrière Hills Volcano, Montserrat. Geochemistry[END_REF][START_REF]Evidence for an Active, Transcrustal Magma System in the Last 60 ka and Eruptive Degassing Budget (H 2 O, CO 2[END_REF][START_REF] Edmonds | Extensive, water-rich magma reservoir beneath southern Montserrat[END_REF][START_REF] Martel | Magma storage conditions and control of eruption regime in silicic volcanoes: experimental evidence from Mt[END_REF][START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF][START_REF] Pichavant | Experimental Parametrization of Magma Mixing: Application to the ad 1530 Eruption of La Soufrière, Guadeloupe (Lesser Antilles)[END_REF][START_REF] Pichavant | Physical conditions, structure, and dynamics of a zoned magma chamber: Mount Pelée (Martinique, Lesser Antilles Arc)[END_REF]Solaro et al., 2019) (Figure II.3b). The proposed mush models for Dominica will be discussed in Part II.2.4.2, but the storage conditions for St. Vincent, Martinique, Guadeloupe and Montserrat are presented here to compare it with Dominica's plumbing system.

For St. Vincent, the storage zone of the magmas is at about 7 km (and up to 3.5 km; 200 MPa; H2O ~4.5-6 wt% and T ~1000 °C), where they evolve prior to eruptions [START_REF] Fedele | Petrological insights on the last 1000 years of explosive activity at La Soufrière volcano, St. Vincent (Lesser Antilles)[END_REF] and pressure ranges determined in glassy inclusions in pumice from the 7 May 1902 eruption and from a similar eruption (1477 cal AD) are between 50 and 450 MPa (depths of in revision). In Martinique, the magmas feeding the latest Plinian and lava dome eruptions of Montagne Pelée originate from a magma storage zone located in pressure ranges of ~50-200 MPa (1.5-6 km) (Balcone-Boissard et al., in revision), 875-900 °C (water contents 5.3-6.3 wt% H2O) [START_REF] Annen | Conditions for the growth of a longlived shallow crustal magma chamber below Mount Pelee volcano (Martinique, Lesser Antilles Arc)[END_REF][START_REF] Martel | Magma storage conditions and control of eruption regime in silicic volcanoes: experimental evidence from Mt[END_REF][START_REF] Pichavant | Physical conditions, structure, and dynamics of a zoned magma chamber: Mount Pelée (Martinique, Lesser Antilles Arc)[END_REF]. More mafic magmas are stored at pressures between 400-500 MPa (14.6-16.2 km below the volcano) and feed the more superficial storage zone (temperature of 1050 °C), such as the magmatic systems in Dominica or Montserrat [START_REF] Boudon | Volcanological evolution of Montagne Pelée (Martinique): A textbook case of alternating Plinian and dome-forming eruptions[END_REF][START_REF] Pichavant | Physical conditions, structure, and dynamics of a zoned magma chamber: Mount Pelée (Martinique, Lesser Antilles Arc)[END_REF]. crust, ranging from 4-16 km (mainly 8-12 km based on petrological studies and 5.5-7.5 based on seismic data; 6-9 wt% H2O).

Most of these systems are then characterized by an extensive mush system, fed by basalt injections from deeper areas that can reactivate the system and enter an unrest phase or lead to an eruption [START_REF] Boudon | Volcanological evolution of Montagne Pelée (Martinique): A textbook case of alternating Plinian and dome-forming eruptions[END_REF][START_REF] Fedele | Petrological insights on the last 1000 years of explosive activity at La Soufrière volcano, St. Vincent (Lesser Antilles)[END_REF][START_REF] Metcalfe | Insights into the Origin and Dynamics of Past Eruptions of La Soufrière de Guadeloupe from Melt Inclusions: Implications for Future Unrest and Environmental Impact[END_REF][START_REF] Murphy | The role of magma mixing in triggering the current eruption at the Soufriere Hills Volcano, Montserrat, West Indies[END_REF][START_REF] Pichavant | Physical conditions, structure, and dynamics of a zoned magma chamber: Mount Pelée (Martinique, Lesser Antilles Arc)[END_REF][START_REF] Pichavant | Crystallization of primitive basaltic magmas at crustal pressures and genesis of the calc-alkaline igneous suite: experimental evidence from St Vincent, Lesser Antilles arc[END_REF]Wadge et al., 2014). In addition, as a result of magmatic differentiation at greater depths (13)(14)(15)(16) km or even 25 km, depending on the dip of the plunging plate and the thickness of the crust), the evolved fluids from this differentiation can rise within the mush and lead to a volcanic eruption [START_REF] Edmonds | Extensive, water-rich magma reservoir beneath southern Montserrat[END_REF].

II.2.4. Dominica island

The island of Dominica belongs to the recent northern arc and has a size of 750 km 2 . This island has the particularity of having the largest number of active volcanoes in the arc (four main volcanoes, compared to one active volcano in general for the other islands of the arc), representing a large volcanic hazard (Lindsay et al., 2005c): Morne aux Diables, Morne Diablotins (1,421 m), Morne Trois Pitons-Micotrin (Morne Trois Pitons: 1,394 m), and the Plats Pays Volcanic Complex (PPVC), from north to south (Figure II.5a-b). These volcanoes mainly produce dome-forming and Plinian eruptions (Lindsay et al., 2005c).

The island volcanic complexes can be divided into three groups according to age and morphology: Miocene , Pliocene , and Pleistocene-recent (2 Ma -present), with the last group including four main potentially active volcanoes (some are satellite lava domes of Morne Trois Pitons-Micotrin). Five large-volume ignimbrite (voluminous pumiceous pyroclastic density current) deposits are now identified as Grande Savane and Layou from Morne Diablotins, Roseau, Grand Fond and Grand Bay from Morne Trois Pitons-Micotrin volcano (Boudon et al., 2017) (Figure II.5a). The oldest volcanic rocks are highly eroded basalts on the east coast [START_REF] Bellon | Reconnaissance chronologique des deux premières phases d'activité volcanique en Dominique (Petites Antilles)[END_REF]Lindsay et al., 2005c;[START_REF] Monjaret | Le Volcanisme de la Dominique. Données chronologiques[END_REF]. From 5.3 Ma to 3.7 Ma, no evidence of volcanism is found on the island, but the Pliocene age centers are characterized by basaltic stratovolcanoes and basaltic andesites in the northern part of the island, which are now highly eroded (Lindsay et al., 2005c). During the Upper Pliocene , the activity occurred in the northern part of the island with lava dome eruptions and pyroclastic density current deposits, from Morne aux Diables and Morne Diablotins (Lindsay et al., 2005c). From the Pleistocene to the recent period (< 1.8 Ma), volcanic activity shifted to the central and southern part of the island, forming six complexes of andesitic/dacitic composition: Morne Trois Pitons (with different satellite volcanoes: Wotten-Waven Micotrin, Morne Watt, Grand Soufrière Hills, Morne Anglais) and PPVC (Lindsay et al., 2005c). The PPVC, in the south of the island, is the area of Dominica where the last magmatic activity was recorded [START_REF] Lindsay | Geological history and potential hazards of the late -Pleistocene to Recent Plat Pays volcanic complex, Dominica, Lesser Antilles[END_REF][START_REF] Roobol | Calderas or gravity-slide structures in the Lesser Antilles island arc?[END_REF]. This complex was partially destroyed by a series of flank collapses where debris avalanches went into the Caribbean Sea [START_REF] Boudon | Volcano flank instability in the Lesser Antilles Arc: Diversity of scale, processes, and temporal recurrence[END_REF][START_REF] Friant | L'île de la Dominique, à l'origine des avalanches de débris les plus volumineuses de l'arc des Petites Antilles[END_REF]. This region is also one of the most active regions of Dominica in terms of seismicity and hot spring activity (1998)(1999)(2000), including on PPVC the Soufrière hot springs and a very active geothermal system on Morne Trois-Pitons Micotrin, at the Valley of Desolation and Boiling Lake [START_REF] Brown | Gas Geochemistry of the Volcanic Hydrothermal Systems of Dominica and St. Lucia, Lesser Antilles: Implications for Volcanic Monitoring[END_REF]Lindsay et al., 2005c;[START_REF] Mayer | Phreatic activity and hydrothermal alteration in the Valley of Desolation, Dominica, Lesser Antilles[END_REF][START_REF] Robson | Some heat measurements in West Indian Soufrières[END_REF][START_REF] Sapper | Ein Besuch von Dominica[END_REF][START_REF] Traineau | Recent geological, geochemical and geophysical surveys of the Roseau Valley, high temperature geothermal field in Dominica,West Indies[END_REF]. The Boiling Lake is the second largest bubbling lake in the world (85x75 m and 10-15 m depth and temperatures between 80 and 90 °C in the last 100 years), the largest being Frying Pan Lake in New Zealand (Figure II.5b) [START_REF] Brown | Gas Geochemistry of the Volcanic Hydrothermal Systems of Dominica and St. Lucia, Lesser Antilles: Implications for Volcanic Monitoring[END_REF]Lindsay et al., 2005c;[START_REF] Mayer | Phreatic activity and hydrothermal alteration in the Valley of Desolation, Dominica, Lesser Antilles[END_REF][START_REF] Robson | Some heat measurements in West Indian Soufrières[END_REF][START_REF] Sapper | Ein Besuch von Dominica[END_REF][START_REF] Traineau | Recent geological, geochemical and geophysical surveys of the Roseau Valley, high temperature geothermal field in Dominica,West Indies[END_REF]. Two historical phreatic eruptions have occurred in this area (in 1880 and 1997) (Lindsay et al., 2005b;[START_REF] Mayer | Phreatic activity and hydrothermal alteration in the Valley of Desolation, Dominica, Lesser Antilles[END_REF]. Decompression experiments and analytical modelling on hydrothermally altered samples of the 1997 phreatic samples permitted to estimate the conditions leading to such events (0.3-0.5 MPa; 115-125 °C) [START_REF] Mayer | Phreatic activity and hydrothermal alteration in the Valley of Desolation, Dominica, Lesser Antilles[END_REF].

Quaternary activity on Dominica consists in the emission of magmas of intermediate to felsic composition in the form of large volume ignimbritic deposits and lava domes (Boudon et al., 2017;Lindsay et al., 2005c;[START_REF] Roobol | Calderas or gravity-slide structures in the Lesser Antilles island arc?[END_REF]Smith et al., 2013). Thus, the island of Dominica has the largest magmatic production in the entire Caribbean arc, which is at least an order of magnitude larger than the other islands (~40 km 3 for the last 100 ka compared to 5-8 km 3 for the neighboring islands of Martinique and Guadeloupe) (Macdonald et al., 2000;Wadge, 1984). This production could be explained by the position of the island relative to the subduction zone and the presence and distribution of a fault system around the island [START_REF] Feuillet | Tectonic context of moderate to large historical earthquakes in the Lesser Antilles and mechanical coupling with volcanoes[END_REF]. Morne Trois Pitons-Micotrin is a large volcanic structure whose present-day summit is composed of three lava domes (with a volume of about 1.5 km 3 ) and the two coalesced lava domes of Micotrin volcano (Figure II.6a) [START_REF] Wills | The Geological History of Southern Dominica and Plutonic Nodules from the Lesser Antilles[END_REF]. The volcanic center of Morne Diablotins is the highest and widest on the island (22 x 18 km, 1421 m high) and is composed of eroded lava domes and deeply gullied pyroclastic covers (Figure II.6b).

Morne Trois Piton-Micotrin volcano produced three ignimbric eruptions (Grand Bay >60 ka, Roseau 33 ka and Grand Fond 24 ka; Volcanic Explosivity Index of 5 (Boudon et al., 2017;Newhall and Self, 1982)), while the Morne Diablotins volcano produced two (Grande Savane and Layou, the latter dated at 50.6 ka), each of which emitted a volume of magma of the order of 5 km 3 DRE. These information and following explanations are based on a detailed study of the eruptions produced during the last 70,000 years of Morne Trois Piton-Micotrin and Morne Diablotins volcanoes, in the central zone of the island by 14 C dating (Boudon et al., 2017) (Figure II.7). These eruptions begin with a Plinian phase generating a convective eruptive column producing pumice and ash fallout (22 to 70 cm thick at ~5 km from the point of emission) and continuing with the emplacement of large volumes of pyroclastic density currents emitted for the most part by overflowing. These deposits are more or less welded and are of several tens of meters thick in the valleys. On the Morne Trois Pitons-Micotrin volcano, a series of eight Plinian eruptions, involving magma volumes at least an order of magnitude smaller (0.5 km 3 DRE), occurred mainly after the last ignimbritic eruption of Grand Fond (VEI of 4) (Boudon et al., 2017;Newhall and Self, 1982). These eruptions generated pumice fallout (~1-2 m thick) and for some of them small-volume of pumiceous pyroclastic density currents resulting from the collapse of the eruptive column. A large part of these deposits is located in the Roseau Valley on the western flank of the volcano (Figure II.5). They have been dated with the 14 C method on carbonized wood contained in the deposits or paleosols. The work in this thesis focused on five eruptions of the Morne Trois Pitons-Micotrin volcano: Grand Fond ignimbritic eruption, Goodwill Entrance: 18 ka cal BP, "Plinian Post Roseau" 1 (PPR1): 12 ka cal BP, PPR2: 12 ka cal BP and PPR3: 9 ka cal BP (Boudon et al., 2017) (Figure II.7). Boudon et al. (2017) showed that the whole rock compositions of the magmas of the studied Plinian eruptions are at the limit of acid andesites and dacites (SiO2: 62-63 wt%; Na2O+ K2O: 4.5-5.5 wt%), comparable to those of the magmas of the ignimbritic eruptions of the same volcano of Morne Trois Pitons-Micotrin volcano. Magmas from ignimbritic eruptions of Morne Diablotins volcano are slightly more differentiated (in whole rock) (Boudon et al., 2017). Residual glasses compositions are comparable in composition between Morne Diablotins and Morne Trois Pitons-Micotrin volcano (Balcone-Boissard et al., 2018). of a heating event on the order of ~25-50 °C about 10 years prior to the eruption [START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF]. Contrary to Halama et al. (2016), that proposed a model of formation of the PPVC lavas by injection and magma mixing with mafic magmas in a poorly evolved reservoir, for these eruptions, there would not have been magma mixing with other more mafic magmas but mixing between different parts of the mush [START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF]. For magmas from the smaller volume Plinian eruptions, they were stored in the upper crust given the pressures determined from their pre-eruptive volatile element contents MPa, ~2-8 km with an average rock density of 2.45 g/cm 3 or 1.8-7.3 km for 2.8 g/cm 3 , 860-895 °C and with 3.25 wt% H2O) [START_REF]Evidence for an Active, Transcrustal Magma System in the Last 60 ka and Eruptive Degassing Budget (H 2 O, CO 2[END_REF]. Comparable volatile element ratios, as well as mineralogical and chemical compositional similarities, may indicate a common source of the magmas that caused the last Grand Fond ignimbritic eruption and the Plinian eruptions, highlighting the presence of a 2-16 km transcrustal mush system (d'Augustin et al., 2020) (Figure II.4b). The pressure and temperature ranges found for the magmas at the origin of the Plinian eruptions of Morne Trois-Pitons Micotrin Plinian agree with other geochemical and experimental petrology works on PPVC [START_REF] Gurenko | A Melt Inclusion Record of Volatiles, Trace Elements and Li-B Isotope Variations in a Single Magma System from the Plat Pays Volcanic Complex, Dominica, Lesser Antilles[END_REF][START_REF] Halama | Preeruptive crystallization conditions of mafic and silicic magmas at the Plat Pays volcanic complex, Dominica (Lesser Antilles)[END_REF]Lindsay et al., 2005c). Comparing these previously found depths for the Morne Trois-Pitons Micotrin magmas for ignimbrites, they are consistent with the boundary between the lower and middle crust of the Lesser Antilles arc crustal structure models (Figure II.4a) [START_REF] Kopp | Deep structure of the central Lesser Antilles Island Arc: Relevance for the formation of continental crust[END_REF][START_REF] Melekhova | Lateral variation in crustal structure along the Lesser Antilles arc from petrology of crustal xenoliths and seismic receiver functions[END_REF]. For magmas at shallower depths (magmas that generated the Plinians eruptions and PPVC ones), these magmas are thought to lie at the boundary between the middle and upper crust (Figure II.4) [START_REF]Evidence for an Active, Transcrustal Magma System in the Last 60 ka and Eruptive Degassing Budget (H 2 O, CO 2[END_REF][START_REF] Gurenko | A Melt Inclusion Record of Volatiles, Trace Elements and Li-B Isotope Variations in a Single Magma System from the Plat Pays Volcanic Complex, Dominica, Lesser Antilles[END_REF][START_REF] Halama | Preeruptive crystallization conditions of mafic and silicic magmas at the Plat Pays volcanic complex, Dominica (Lesser Antilles)[END_REF]Lindsay et al., 2005c). As most of the islands of the Lesser Antilles arc, Dominica shows a well-developed transcrustal magmatic system and the discontinuity between the upper and the middle crust appears to be a major magma ponding zone beneath most of the islands (Figure

. Tectonic and structural context of Kamchatka

The Kamchatka region is part of the "Pacific Ring of Fire", which is a long chain of 40,000 km long with more than 450 active volcanoes producing explosive and effusive eruptions [START_REF] Eichelberger | Preface, in: Volcanism and Subduction: The Kamchatka Region[END_REF] (Figure II.8a). This region forms a subduction cusp junction within the 6000 km long Kuriles-Kamchatka-Aleutian-Alaska subduction, where the Kamchatka and Aleutian Islands arc systems meet perpendicularly at the Kamchatka Peninsula (Figure II.8) [START_REF] Eichelberger | Introduction: Subduction's sharpest arrow[END_REF][START_REF] Scholl | Viewing the tectonic evolution of the Kamchatka-Aleutian (KAT) connection with an Alaska crustal extrusion perspective[END_REF]. The chain of the Emperor Seamounts is at the junction between the two arc systems. The Kamchatka-Aleutian-Alaska subduction segment consists of two parts: the Kamchatka mainland peninsula with the Kurile oceanic Islands (archipelago of about 30 islands with about 100 volcanoes of which 40 are active) marked by a northeastward subduction zone and that of the Alaska Peninsula with the Aleutian Islands, an archipelago of 300 volcanic islands, where the plate boundary becomes a northwestward transform fault [START_REF] Scholl | Viewing the tectonic evolution of the Kamchatka-Aleutian (KAT) connection with an Alaska crustal extrusion perspective[END_REF] (Gorbatov et al., 1997;[START_REF] Kayzar | Deciphering petrogenic processes using Pb isotope ratios from time-series samples at Bezymianny and Klyuchevskoy volcanoes, Central Kamchatka Depression[END_REF][START_REF] Konstantinovskaia | Arc-continent collision and subduction reversal in the Cenozoic evolution of the Northwest Pacific: an example from Kamchatka (NE Russia)[END_REF][START_REF] Koulakov | Slab shape in subduction zones beneath the Kurile-Kamchatka and Aleutian arcs based on regional tomography results[END_REF]Minster and Jordan, 1978;[START_REF] Watson | Tectonic evolution of Kamchatka and the sea of Okhotsk and implications for the Pacific Basin[END_REF]. The Pacific plate entering the Kamchatka trench has an inferred age of ~95 Ma at the southern tip of Kamchatka (Portnyagin et al., 2007a;[START_REF] Renkin | Depth and age in the north Pacific[END_REF]. For some authors, Kamchatka and eastern Siberia constitute the western end of the Northern American Plate Copyrighted material, unauthorised distribution (Figure 1; [START_REF] Koulakov | Slab shape in subduction zones beneath the Kurile-Kamchatka and Aleutian arcs based on regional tomography results[END_REF][START_REF] Demets | Oblique convergence and deformation along the Kuril and Japan Trenches[END_REF][START_REF] Fujita | Seismicity and focal mechanisms of the Arctic region and the North American plate boundary in Asia[END_REF][START_REF] Kogan | Geodetic constraints on the rigidity and relative motion of Eurasia and North America[END_REF]Levin et al., 2002;[START_REF] Park | A Dangling Slab, Amplified Arc Volcanism, Mantle Flow and Seismic Anisotropy in the Kamchatka Plate Corner[END_REF] although others believe that they are part of the Okhotsk microplate [START_REF] Riegel | Extrusion tectonics of the Okhotsk Plate, northeast Asia[END_REF][START_REF] Zonenshain | Introduction to geodynamics[END_REF]. Three significant features of the Kamchatka segment make it different from the southern part of the peninsula. First of all, three volcanic zones are present in Kamchatka: the Sredinny Range (Central-Kamchatka belt) and an Eastern zone subdivided in two: the Central Kamchatka Depression (CKD) and the Eastern Volcanic Front, where all the historical eruptions occurred [START_REF] Ponomareva | Late Pleistocene-Holocene volcanism on the Kamchatka Peninsula, Northwest Pacific Region, in: Volcanism and Subduction: The Kamchatka Region[END_REF]. The Eastern zone is much younger than the Kuriles volcanic segment. Furthermore, the eastern coast of the Kamchatka segment is composed of three small peninsulas (Shipunski, Kronotski and Kamchatkaski Mys) constituted by Paleocene-Eocene volcanic and sedimentary series (see 1-3 in Figure II.9a). These peninsulas are considered as fragments of an ancient island arc, also referred as the "Kronotski arc". Finally, the seismic zone is relatively shallow, the depth of the slab in the modern Kamchatka subduction zone is at least 300 km less compared to the depth of the Kuriles slab (at more than 800 km) [START_REF] Lander | The origin of the modern Kamchatka Subduction Zone[END_REF] Unlike South Kurile Islands, which formed during the Oligocene, the modern Kamchatka structure formed only during the last few million years [START_REF] Lander | The origin of the modern Kamchatka Subduction Zone[END_REF]. The Central-Kamchatka volcanic belt, that goes along the Sredinny range, formed at that time but has now almost lost all activity (except from two volcanoes) but was still active during the Holocene (Figure II.9a).

After the end of the Miocene (7 to 10 Ma), a collision of an arc called the "the Kronotski arc" (Figure (Gorbatov et al., 1997;[START_REF] Lander | The origin of the modern Kamchatka Subduction Zone[END_REF][START_REF] Peyton | Mantle flow at a slab edge: Seismic anisotropy in the Kamchatka Region[END_REF][START_REF] Portnyagin | Mantle temperature control on composition of arc magmas along the Central Kamchatka Depression[END_REF]. The Central Kamchatka Depression was then developed as a fore-arc basin associated with the former subduction below the Sredinny Range [START_REF] Koulakov | Mantle and Crustal Sources of Magmatic Activity of Klyuchevskoy and Surrounding Volcanoes in Kamchatka Inferred From Earthquake Tomography[END_REF]

(Figure II.9a).
Results of seismic tomography as well as geophysical and geological data show that the upper mantle beneath Kamchatka has undergone major changes since 10 Ma and that subduction dynamics beneath Kamchatka are not stable (Levin et al., 2002). According to these results, at least two episodes of slab detachment have occurred: a first one between 5 and 10 Ma, where part of the slab created by subduction of the Komandorsky Basin lithosphere detached and a second one since 2 Ma, where a fragment of the present-day Pacific Plate detached and sank into the mantle (Figure II.9c). Following these slab detachments events, warmer asthenospheric material flowed around the subducting slab and upwards towards the surface (Levin et al., 2002) (Figure II.9c). These slab detachment episodes are associated to different volcanic events such as a progressive change from island-arc volcanism to rift-like volcanism at about 5 Myr ago, suggesting thermal rejuvenation of the mantle lithosphere beneath Kamchatka since that time [START_REF] Fedorov | Neogene volcanics of the Kamchatka Isthmus and geodynamics of the Aleutian-Kamchatka Junction[END_REF]Levin et al., 2002;[START_REF] Volynets | Geochemical Types, Petrology, and Genesis of Late Cenozoic Volcanic Rocks from the Kurile-Kamchatka Island-Arc System[END_REF]. Active volcanism in the Klyuchevskoy group lies above the slab edge itself and dates back only from 50 ky ago (Levin et al., 2002).

Regarding the historical volcanic activity in Kamchatka, numerous volcanic eruptions occurred at the Kamchatka segment (as well as some of the ten largest earthquakes of the 20th century and tsunamis) [START_REF] Eichelberger | Preface, in: Volcanism and Subduction: The Kamchatka Region[END_REF]. These include the flank collapse events of Bezymianny and Shiveluch in 1956 and 1964, respectively, and the 1975 eruption of Tolbachik, referred to as the "Great Tolbachik Fissure Eruption" with a volcanic plume that went as far as 30 km into the atmosphere [START_REF] Eichelberger | Introduction: Subduction's sharpest arrow[END_REF]. As of today, 4 volcanoes are erupting or are in the awakening stage in the Kamchatka system: Klyuchevskoy (February 2021-March 2021), Shiveluch (January 2019-December 2021 and still erupting), Karimsky (March 2021-now), and Bezymianny (August 2022) (Kamchatka Volcanic Eruption Response Team (KVERT), 2022). These areas are monitored by different observatories to study natural hazards and seismic, volcanic and tsunami activity in the North Pacific basin that could threaten people (as one hundred volcanoes on this subduction segment would be capable of generating ash clouds at airline altitudes, representing risks to air traffic) [START_REF] Eichelberger | Preface, in: Volcanism and Subduction: The Kamchatka Region[END_REF]. Volcanic hazard estimation has been done on many Holocene volcanoes [START_REF] Bazanova | Expected hazard from the Avachinsky volcano eruptions, in: Geodynamics and Volcanism of the Kurile-Kamchatka Island Arc System[END_REF][START_REF] Melekestsev | Prediction of volcanic hazards on the basis of the study of dynamics of volcanic activity, Kamchatka[END_REF][START_REF] Ponomareva | Volcanic Hazards Assessment in the Area of lake Kronotskoye, Uzon caldera and Valley of Geysers[END_REF], 80% of Kamchatka's inhabitants are concentrated around three major cities: Petropavlovsk-Kamchatsky and Elizovo, ~30 km south of the Koriaksky and group appear to be predominantly calc-alkaline but may exhibit abrupt changes [START_REF] Koulakov | Rapid changes in magma storage beneath the Klyuchevskoy group of volcanoes inferred from time-dependent seismic tomography[END_REF].

It includes Klyuchevskoy volcano, the most active arc volcano in the world [START_REF] Fedotov | Active volcanoes of Kamchatka[END_REF][START_REF] Kersting | Klyuchevskoy Volcano, Kamchatka, Russia: The Role of High-Flux Recharged, Tapped, and Fractionated Magma Chamber(s) in the Genesis of High-Al2O3 from High-MgO Basalt[END_REF][START_REF] Melekestsev | Volcanism and Relief Formation[END_REF]Portnyagin et al., 2007a) [START_REF] Dobretsov | Migration paths of magma and fluids and lava compositions in Kamchatka[END_REF]. Another cluster of volcanoes is found in the CKD, the Shiveluch Volcanic Group which includes the Shiveluch, Zarechny and Kharchinsky volcanoes [START_REF] Ponomareva | Late Pleistocene-Holocene volcanism on the Kamchatka Peninsula, Northwest Pacific Region. in Volcanism and Subduction[END_REF]Portnyagin et al., 2007a;Volynets et al., 1999b) (Figure II.9a;Figure II.11). A body of slow material below the Komandorsky basin may be responsible for a "slab melt" or adakite, petrological signature at Shiveluch, through partial melting of the slab edge [START_REF] Yogodzinski | Geochemical evidence for the melting of subducting oceanic lithosphere at plate edges[END_REF]. Further north along the CKD, another volcanic cluster is the Shisheisky volcanic complex and is located ~50 km north of Shiveluch volcano and may have been formed above the slab edge (Portnyagin et al., 2007a). Finally, at the northern end of the CKD, two volcanoes are present: Nachikinsky and Khailulia [START_REF] Portnyagin | Transition from arc to oceanic magmatism at the Kamchatka-Aleutian junction[END_REF] and no subducting panels are present beneath these volcanoes (Portnyagin et al., 2007a). From the south to the north of the CKD, a decrease in the size of volcanic structures and the rate of magmatic production during the Quaternary are observed [START_REF] Fedotov | Active volcanoes of Kamchatka[END_REF][START_REF] Ponomareva | Late Pleistocene-Holocene volcanism on the Kamchatka Peninsula, Northwest Pacific Region, in: Volcanism and Subduction: The Kamchatka Region[END_REF]Portnyagin et al., 2007a).

A recent model from [START_REF] Koulakov | Mantle and Crustal Sources of Magmatic Activity of Klyuchevskoy and Surrounding Volcanoes in Kamchatka Inferred From Earthquake Tomography[END_REF] using a tomographic model, explained the formation of the western and eastern segments of the CKD (Figure II.13). After the first slab detachment (Figure II.10a3) and the beginning of the new subduction, asthenosphere uplift leads to volcanism in the center of the CKD (Ushkovsky and Shiveluch) and active rifting. Following this phase to the present time, Eastern CKD is formed as a rift basin and volcanic activity occurs along the Tolbachik-Bezymianny-Klyuchesvkoy fault zone [START_REF] Koulakov | Mantle and Crustal Sources of Magmatic Activity of Klyuchevskoy and Surrounding Volcanoes in Kamchatka Inferred From Earthquake Tomography[END_REF].

II.3.2. Magma source and compositions

Seismic data show smaller dips of the dipping panel under the CKD, ranging from 45° below the Tolbachik to 25° below the Shiveluch [START_REF] Fedotov | Active volcanoes of Kamchatka[END_REF]Gorbatov et al., 1997;Park et al., 2002;Portnyagin et al., 2007a). Upward flexure of the Pacific plate toward the Kamchatka-Aleutian junction results in large variations in subducting slab depths (80-180 km) beneath active CKD volcanoes while the distance to the Kamchatka volcanoes remains constant (200-250 km), implying a decrease in mantle wedge thickness as the crustal thickness remains around 30-35 km (Balesta, 1991;[START_REF] Bogdanov | Tectonic map of the Sea of Okhotsk region: map and explanatory notes[END_REF][START_REF] Park | A Dangling Slab, Amplified Arc Volcanism, Mantle Flow and Seismic Anisotropy in the Kamchatka Plate Corner[END_REF]Portnyagin et al., 2007a). This length of mantle column can be potentially involved in magma generation and decreases from ~165 km beneath Tolbachik to ~55 km beneath Shiveluch (Portnyagin et al., 2007a).

The significant volcanic activity in the CKD may then be due to intra-arc rifting and magmatic upwelling, generating partial melting by decompression as well as significant flow of fluids from the Emperors submarine chain. Alkaline magmas are produced by volcanoes at the north of the CKD, as lavas from Klyuchevskoy volcano that have characteristics of high equilibrium temperatures [START_REF] Kersting | Klyuchevskoy Volcano, Kamchatka, Russia: The Role of High-Flux Recharged, Tapped, and Fractionated Magma Chamber(s) in the Genesis of High-Al2O3 from High-MgO Basalt[END_REF]Levin et al., 2002;[START_REF] Ponomareva | Late Pleistocene-Holocene volcanism on the Kamchatka Peninsula, Northwest Pacific Region, in: Volcanism and Subduction: The Kamchatka Region[END_REF]. The upwelling could be caused by detachment of lithospheric material (Portnyagin et al., 2007a). The rocks of the southern CKD are calcalkaline basalt-andesite medium-Ks (Ponomareva et al., 2007) (Figure II.12a). Those of the northern CKD have different trace element concentrations typical of adakites [START_REF] Defant | Derivation of some modern arc magmas by melting of young subducted lithosphere[END_REF]. A geochemical zonation along this arc can therefore be observed, from south to north, with a subduction signature in the south to an intraplate/oceanic signature in the north (Figure II.12a). Variations in CKD magmas also result from several processes including primary magmas drifting from the MORB source (following fluid flow from the subducting plate), small amounts of sediment-derived fluids or liquids being added to these magma-generating zones, as well as partial melting of the Pacific plate beneath Tolbachik volcano, Klyuchevskoy volcano, Shiveluch volcano, Shisheisky Complex and Nachikinsky volcano). Grey squares represent analyses of CKD volcanic rocks not denoted by other symbols. b1-b2) Possible scenarios for magma generation beneath the CKD. b1) Mixing model that suggests that hydrous eclogitic melt triggers mantle melting. Eclogitic melts react with mantle peridotite, producing high-Mg# pyroxenites in the mantle wedge that are melted later with surrounding peridotite. b2) Reaction model suggests interactions between eclogitic melts and mantle-wedge peridotite, with mantle melting triggered by slab melt. Eclogitic melts leave the slab and interact with the mantle wedge on their rise according to the AFC (Assimilation-Fractional Crystallization) or flux mantle melting scenario (assimilation of mantle material by the rising melts as the temperature of these melts is high). Shallow assimilation may follow mantle melting and can occur in the overriding plate (Portnyagin et al., 2007a).

The tomographic model of [START_REF] Koulakov | Mantle and Crustal Sources of Magmatic Activity of Klyuchevskoy and Surrounding Volcanoes in Kamchatka Inferred From Earthquake Tomography[END_REF] shows that the volcanism of the Klyuchevskoy group and volcanoes in the north of the CKD (such as Shiveluch) developed after the major subduction reconfiguration following the Miocene-Pliocene collisions of the Kronotsky arc terrane (Figure II.10a3) and has been fed by two sources: the rifting caused by the asthenospheric flow from the Aleutian-Kamchatka slab window and the upwelling of the volatile-rich fluids from the slab, typical for subduction zones (Figure II.13). The exceptional level of volcanic activity and the diversity of volcanic manifestations in this region is likely the result of these two sources, especially the large number of basaltic shield volcanoes not typical for subduction volcanic chains [START_REF] Koulakov | Mantle and Crustal Sources of Magmatic Activity of Klyuchevskoy and Surrounding Volcanoes in Kamchatka Inferred From Earthquake Tomography[END_REF].

II.3.3. Structure of the crust

The particularly intense volcanic activity in this region can be due to the hot mantle material brought by the around-slab-edge asthenospheric flow (Levin et al., 2002) and related crustal extension [START_REF] Koulakov | Mantle and Crustal Sources of Magmatic Activity of Klyuchevskoy and Surrounding Volcanoes in Kamchatka Inferred From Earthquake Tomography[END_REF]. The area of the northern group of volcanoes of the CKD is covered by 25 telemetered seismic stations of the Kamchatka Branch of Geophysical Survey (KBGS) which permits a well study of this area, from the uppermost crust (0-10 km), lower crust (15-35 km) to the mantle (> 35 km) [START_REF] Koulakov | Mantle and Crustal Sources of Magmatic Activity of Klyuchevskoy and Surrounding Volcanoes in Kamchatka Inferred From Earthquake Tomography[END_REF]. In the crust, large hidden parts of shield volcanoes can be seen beneath the CKD (P and S-wave high velocity anomalies beneath Krestovsky and Ushkovsky) and high velocity anomalies in the upper crust mark the presence of large igneous bodies beneath other volcanoes (Shiveluch, Tolbachik, Nibolka and Kizimen) [START_REF] Koulakov | Mantle and Crustal Sources of Magmatic Activity of Klyuchevskoy and Surrounding Volcanoes in Kamchatka Inferred From Earthquake Tomography[END_REF]. Nearly continuous occurrence of deep long-period earthquakes near the crust-mantle boundary (Melnik et al., 2020;[START_REF] Shapiro | Deep and shallow long-period volcanic seismicity linked by fluid-pressure transfer[END_REF] is likely related to the underplating of mafic magmas and widespread seismic tremors mark an active trans-crustal magmatic system connecting different volcanoes of this group [START_REF] Journeau | Seismic tremor reveals active trans-crustal magmatic system beneath Kamchatka volcanoes[END_REF]. Well-developed magma chambers beneath some of most active volcanoes such as Bezymianny are well documented, with the first high-resolution structures being identified beneath the Tolbachik volcanic system [START_REF] Koulakov | Three different types of plumbing system beneath the neighboring active volcanoes of Tolbachik, Bezymianny, and Klyuchevskoy in Kamchatka[END_REF] (Figure II.13a). The three volcanoes in the Klyuchevskoy group (Klyuchevskoy, Bezymianny and Tolbachik) have completely different compositions and eruptive styles. The tomographic study of [START_REF] Koulakov | Three different types of plumbing system beneath the neighboring active volcanoes of Tolbachik, Bezymianny, and Klyuchevskoy in Kamchatka[END_REF] showed three different types of feeding system for these volcanoes: 1) the basaltic lavas of Klyuchevskoy are formed from a reservoir at 25-30 km through a nearly vertical pipe-shaped conduit; 2) the explosive Bezymianny volcano is supplied through a dispersed system of crustal reservoirs where a lighter felsic material separates from the mafic component at depth and ascends to the upper crust to form andesitic magma sources (< 10 km); 3) for Tolbachik, low-viscosity volatile-saturated basalts ascend from two deep reservoirs following a system of fractures in the crust (dike complexes) associated with the intersections of regional faults [START_REF] Koulakov | Three different types of plumbing system beneath the neighboring active volcanoes of Tolbachik, Bezymianny, and Klyuchevskoy in Kamchatka[END_REF] The eruptive events following a collapse (or a rockslide) of part of a volcano depend on the factors that caused it: in most cases, when magma rises from depth, the cone of a volcano loses its stability and rockslide occurs as magma is still deep enough [START_REF] Belousov | The Shiveluch volcanic eruption of 12 November 1964-explosive eruption provoked by failure of the edifice[END_REF]. But, when magma rises close to the surface as a lava dome or cryptodome and a rockslide occurs, the scenario of a catastrophic lateral blast immediately follows the flank collapse [START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF]. This explains why lateral blasts are rather rare.

The 1955-1956 eruption was immediately followed by the quasi-continuous growth of lava domes, with extrusive-explosive behavior, until 1976 (Figure II.16f). From 1977 to the present, extrusive-explosive and effusive periods followed, with phases of more or less important extrusion and phases of destruction by sector collapse, generating ash clouds and many ash and block pyroclastic density currents, which were often followed by effusions of lava flows [START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF]Turner et al., 2013) (Figure II.15d;Figure II.16g). These short-lived strong Vulcanian explosive eruptions occurred from the lava dome summit once or twice a year, with ash clouds to heights up to 8-15 km (a.s.l.), and with pyroclastic density currents extending about 12-13 km away from the lava dome [START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]. Only few eruptions during the 1980s and 1990s were characterized by effusive activity or lava flows prior to explosions [START_REF] Belousov | Pyroclastic surges and flows from the 8-10 May 1997 explosive eruption of Bezymianny volcano, Kamchatka, Russia[END_REF][START_REF] Mania | Deformations and Morphology Changes Associated with the 2016-2017 Eruption Sequence at Bezymianny Volcano, Kamchatka[END_REF] Thus, over its eruptive activity, Bezymianny volcano has shown a diversity of eruption styles, with a duality between Vulcanian eruptions and lava dome growth eruptions that can generate flank-collapse and powerful laterally directed explosions [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]. Nowadays, Bezymianny old eruptions remnants are visible in its morphology, with the amphitheater related to the 1956 flank collapse and a 500-600 m high lava dome in the center [START_REF] Mania | Deformations and Morphology Changes Associated with the 2016-2017 Eruption Sequence at Bezymianny Volcano, Kamchatka[END_REF]

II.3.4.2. Plumbing system

Bezymianny lavas are andesites with rare basalts, basaltic andesites and dacites from the calc-alkaline series [START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF]. [START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF] showed that whole rock concentrations of Bezymianny products were becoming more mafic since the eruption of 1956 by the loss of incompatible elements (2.5 wt% MgO and 61 wt% SiO2 in 1956 against 4 wt% MgO and 57 wt% SiO2 in 2010). Bezymianny and Kamen lavas belong to the same differentiation series, with Bezymianny ones being more differentiated (Churikova et al., 2013a;d'Augustin, 2021). This could indicate a common reservoir beneath the two volcanoes or that Bezymianny reservoir is Kamen's old one (Churikova et al., 2013a). Bezymianny plumbing system may root deeply, and it is thought that Bezymianny, Klyuchevskoy, but also Tolbachik, located further south, derive their fluids from a common deep parental magma chamber at 30 km depth [START_REF] Koulakov | Three different types of plumbing system beneath the neighboring active volcanoes of Tolbachik, Bezymianny, and Klyuchevskoy in Kamchatka[END_REF]Shapiro et al., 2017) (Figure II.13). During magma ascent beneath Bezymianny, the volatile-rich magma is stored at different levels [START_REF] Mania | Deformations and Morphology Changes Associated with the 2016-2017 Eruption Sequence at Bezymianny Volcano, Kamchatka[END_REF]. These levels are likely associated with magma chambers, at approximately 15 km and 5 km, but also possibly at 1-1.5 km depth, with a volatile zone suggested by the study of the seismicity of Bezymianny eruption sequence in 2007 [START_REF] Davydova | Petrology of mafic enclaves in the 2006-2012 eruptive products of Bezymianny Volcano, Kamchatka[END_REF][START_REF] Koulakov | Three different types of plumbing system beneath the neighboring active volcanoes of Tolbachik, Bezymianny, and Klyuchevskoy in Kamchatka[END_REF][START_REF] Mania | Inflating Shallow Plumbing System of Bezymianny Volcano, Kamchatka, Studied by InSAR and Seismicity Data Prior to the 20 December 2017[END_REF][START_REF] Mania | Deformations and Morphology Changes Associated with the 2016-2017 Eruption Sequence at Bezymianny Volcano, Kamchatka[END_REF][START_REF] Ozerov | A Petrologic-Geochemical Model for Genetic Affinity between Basaltic and Andesitic Volcanisms at Klyuchevskoi and Bezymyannyi Volcanoes[END_REF]Shcherbakov et al., 2011;[START_REF] Ter Heege | Experimental Determination of Fe-Mg Interdiffusion Coefficients in Orthopyroxene Using Pulsed Laser Ablation and Nanoscale Thin Films[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF] Numerous studies have estimated the storage conditions of the magmas prior to the 1956 climactic activity of Bezymianny, either by compositions of phenocrysts or by experimental petrology (Almeev et al., 2013a(Almeev et al., , 2013b[START_REF] Almeev | Problems of the Stoichiometry and Thermobarometry of Magmatic Amphiboles: An Example of Hornblende from the Andesites of Bezymyannyi Volcano, Eastern Kamchatka[END_REF][START_REF] Kadik | Physicochemical Crystallization Conditions and Genesis of Andesites[END_REF][START_REF] Neill | Influence of pre-eruptive degassing and crystallization on the juvenile products of laterally directed volcanic explosions[END_REF][START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF][START_REF] Shcherbakov | Phase equilibria constraints on preeruptive magma storage conditions for the 1956 eruption of Bezymianny Volcano, Kamchatka, Russia[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF].

Practically all models infer the presence of two magmatic reservoirs at Bezymianny, but with different P-T conditions and thus different estimated depths:

-A last storage between ~2-4 km shortly prior to the 1956 climactic phase (phase equilibria and melt inclusions; 50-100 MPa, 890-930 °C, 2.4 ± 0.9 H2O wt%) and another one deeper, at ~8 km (200 MPa, 775 °C; for a density of 2700 kg.m -3 ) [START_REF] Shcherbakov | Phase equilibria constraints on preeruptive magma storage conditions for the 1956 eruption of Bezymianny Volcano, Kamchatka, Russia[END_REF]). -Deeper: crystallization conditions of phenocrysts were estimated at 11-30 km (300-800 MPa, 875-925 °C and > 6 wt% H2O) [START_REF] Kadik | Physicochemical Crystallization Conditions and Genesis of Andesites[END_REF], agreeing with [START_REF] Almeev | Problems of the Stoichiometry and Thermobarometry of Magmatic Amphiboles: An Example of Hornblende from the Andesites of Bezymyannyi Volcano, Eastern Kamchatka[END_REF], reporting a pressure of ~600 MPa (23 km) and T of 800-920 °C and with [START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF], with amphibole reaction rims development. Phase equilibria simulations on Bezymianny volcanic subseries indicate that parental melts are stored at ~15 km depth under a pressure of 490-520 MPa (1,100 °C, 1.5-2 wt% H2O) (Almeev et al., 2013a) or between 250-700 MPa with geothermobarometry (Almeev et al., 2013b). These two last pressure ranges correspond well to the crustal reservoirs characterized by seismic investigations at ~7 km and 18 km [START_REF] Fedotov | The magmatic system of the Klyuchevskaya group of volcanoes inferred from data on its eruptions, earthquakes, deformation, and deep structure[END_REF][START_REF] Koulakov | Three different types of plumbing system beneath the neighboring active volcanoes of Tolbachik, Bezymianny, and Klyuchevskoy in Kamchatka[END_REF][START_REF] Ter Heege | Experimental Determination of Fe-Mg Interdiffusion Coefficients in Orthopyroxene Using Pulsed Laser Ablation and Nanoscale Thin Films[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF].

The mineral assemblage of the ash generated by the eruptive activity from October to November is different from the one in March 1956 (dominated by orthopyroxene or amphibole, respectively) [START_REF] Gorshkov | Bezymianny Volcano and Peculiarities of Its Last Eruptions in[END_REF][START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF]) so the magmatic reservoir should have experienced significant transformations with a change in the phenocryst assemblage. The magma at the origin of the volcanic activity reported in October-November 1955 is thus unrelated to the magma at the origin of the directed lateral blast. Following the study of rims disequilibria of these amphiboles from 1956 cryptodome fragments, they were developed at a pressure of approximately 600 MPa, following heating of the magma chamber from 890 to 1050 °C in less than 37 days [START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF]. This heating could have been caused by a fresh magma injection in February 1956, generating a significant heating of the magmatic chamber [START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF]. [START_REF] Shcherbakov | Phase equilibria constraints on preeruptive magma storage conditions for the 1956 eruption of Bezymianny Volcano, Kamchatka, Russia[END_REF] showed, following amphibole breakdown rims from the Plinian phase, that magmas from the Plinian phase were stored at 3 km depth prior to the final ascent for at least 40 days.

Another model of the pre-eruptive dynamics showed that the 1956 climactic volcanic activity was fed by a three-component magma mixing system (end-member-mixing-analysis), with two main reservoirs and a mafic injection that occurred prior to eruption [START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF] There is a rapid ascent of magma from this lower reservoir, as seen by the survival of mantle xenoliths, without disintegration [START_REF] Davydova | Petrology of mafic enclaves in the 2006-2012 eruptive products of Bezymianny Volcano, Kamchatka[END_REF]. Deposits contemporaneous with eruptions show evidence of magma mixing, such as mafic enclaves, imbalanced xenocrystal assemblages, and reverse zoned crystals characteristics of injections [START_REF] Davydova | Petrology of mafic enclaves in the 2006-2012 eruptive products of Bezymianny Volcano, Kamchatka[END_REF]Shcherbakov et al., 2011;[START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]. Chronometric diffusion studies have been carried out on this volcano by Davydova et al. (2018), using zoned orthopyroxene crystals from samples of eruptions from 2006 to 2012. These authors showed that for some eruptions, the orthopyroxene rims formed up to 3 years before the onset of the eruption (correlated to high seismic activity of the volcano) and for others, 2 months before (Davydova et al., 2018). These differences are caused by the different paths taken during magmatic recharge to the shallower Bezymianny chamber (Davydova et al., 2018). Explosive activity in December 2017 was also triggered by magma influx from deeper levels of the plumbing system shortly before the eruption [START_REF] Davydova | Petrological evidence of rapid evolution of the magma plumbing system of Bezymianny volcano in Kamchatka before the December 20th, 2017 eruption[END_REF]. II.3.4.3. Monitoring The first detailed seismological observations of Kamchatka started in 1961 (by the Pacific Ocean seismic expedition of the Institute of Physics of the Earth Russian Academy of Sciences and other institutes) [START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF]. In 2000, the Kamchatka Branch of Geophysical Survey, from the Russian Academy of Sciences, started monitoring the volcanic activity in Kamchatka [START_REF] Senyukov | Monitoring of volcanic activity in Kamchatka by remote sensing techniques in 2000-2004[END_REF][START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF].

For the 1955-1956 eruption of Bezymianny, only one seismic station was in operation and was located at 42 km away from the volcano [START_REF] Gorshkov | Bezymianny Volcano and Peculiarities of Its Last Eruptions in[END_REF][START_REF] Tokarev | The prediction of large explosions of andesitic volcanoes[END_REF][START_REF] Tokarev | Volcanic Earthquakes of Kamchatka[END_REF]. The sites of seismic foci were determined roughly but the uninterrupted operation of this well-equipped seismic station allowed to obtain material to study the seismic regime [START_REF] Gorshkov | Bezymianny Volcano and Peculiarities of Its Last Eruptions in[END_REF][START_REF] Tokarev | The prediction of large explosions of andesitic volcanoes[END_REF][START_REF] Tokarev | Volcanic Earthquakes of Kamchatka[END_REF]. A 23-day swarm of earthquakes was detected from 29 September 1955 and by 19 October 1955, it reached the level of hundreds of seismic events per day with hypocenters below Bezymianny (37 thousand volcanic earthquakes recorded from September to October) [START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF][START_REF] Tokarev | The prediction of large explosions of andesitic volcanoes[END_REF][START_REF] Tokarev | Volcanic Earthquakes of Kamchatka[END_REF]. A catalogue of the preceding and accompanying earthquakes is provided by [START_REF] Tokarev | Volcanic Earthquakes of Kamchatka[END_REF]. The first seismic station near the volcano was installed in September 1960 (Apakhonchich, 16 km northeast of Bezymianny) [START_REF] Bogoyavlenskaya | Bezymianny volcano[END_REF]Girina, 2013) (Figure II.18). Nowadays, volcano monitoring allows an identification of the eruption precursors from days to weeks before eruptions, increasing aviation safety over Kamchatka [START_REF] Chebrov | The system of detailed seismological observations in Kamchatka in 2011[END_REF][START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF][START_REF] Mania | Deformations and Morphology Changes Associated with the 2016-2017 Eruption Sequence at Bezymianny Volcano, Kamchatka[END_REF][START_REF] Senyukov | Monitoring and prediction of volcanic activity in Kamchatka from seismological data: 2000-2010[END_REF]. A permanent network of seismic stations is monitoring the volcanoes of the Klyuchevskoy group (18 stations in 2009 and 19 stations from 2010 onward) [START_REF] Soubestre | Depth Migration of Seismovolcanic Tremor Sources Below the Klyuchevskoy Volcanic Group (Kamchatka) Determined From a Network-Based Analysis[END_REF]. Three types of volcanic earthquakes and volcanic tremor are typical of Bezymianny and accompany lava dome growth, lava extrusion at the surface and explosions from the lava dome [START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF][START_REF] Tokarev | Volcanic Earthquakes of Kamchatka[END_REF]. Volcanic tremor is usually related to paroxysmal phase of the eruption [START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF]. Long-period seismicity may also indicate heralding eruptions but earlier studies showed that only one out of four eruptions were preceded by long-period events [START_REF] Mania | Deformations and Morphology Changes Associated with the 2016-2017 Eruption Sequence at Bezymianny Volcano, Kamchatka[END_REF][START_REF] Shapiro | Deep and shallow long-period volcanic seismicity linked by fluid-pressure transfer[END_REF][START_REF] West | Recent eruptions at Bezymianny volcano-A seismological comparison[END_REF].

GPS campaigns prior to eruptions between 2005-2010 revealed minor deformation with a shallow and small reservoir [START_REF] Grapenthin | Surface deformation of Bezymianny Volcano, Kamchatka, recorded by GPS: The eruptions from 2005 to 2010 and long-term, longwavelength subsidence[END_REF][START_REF] Mania | Inflating Shallow Plumbing System of Bezymianny Volcano, Kamchatka, Studied by InSAR and Seismicity Data Prior to the 20 December 2017[END_REF]. Remote sensing techniques are also used for identification of precursory activity (webcam images, high-resolution satellite radar data to study morphology changes, for example), as enhanced ground temperature anomalies are a common precursor for Bezymianny eruptions [START_REF] Carter | ASTER and field observations of the 24 December 2006 eruption of Bezymianny Volcano, Russia[END_REF][START_REF] Mania | Deformations and Morphology Changes Associated with the 2016-2017 Eruption Sequence at Bezymianny Volcano, Kamchatka[END_REF][START_REF] Ramsey | Spaceborne observations of the 2000 Bezymianny, Kamchatka eruption: the integration of high-resolution ASTER data into near real-time monitoring using AVHRR[END_REF][START_REF] Van Manen | Satellite thermal observations of the Bezymianny lava dome 1993-2008: Precursory activity, large explosions, and dome growth[END_REF]. Copyrighted material, unauthorised distribution (Figure 1; [START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF] of the Shchapinsky graben, within a fault system including the junction of the graben with the Tumrok ridge horst, a network of large amplitude north-east striking normal faults (Browne et al., 2010) (Figure II.19). It has three remarkable geological features: this fault system intersects pre-Kizimen deposits and the volcano itself, such as the SW-NE normal fault located at the northern-west (NW) slope of the volcano (Auer et al., 2018). Since 1825, a fumarole has been present continuously at the NW slope of the volcano at 1950 m a.s.l. to 400 m below the volcano summit [START_REF] Kirsanova | Fumarolic activity of Shiveluch and Kizimen volcanoes in 1979-1980[END_REF][START_REF] Piip | Kizimen volcano[END_REF]. Four cycles of eruptive activity were identified at Kizimen, dated from 12-11 ka to the present, prior to the recent eruption in 2010, through a tephrochronological study (KZI to KZIV) (Melekestsev et al., 1995). Each cycle begins with large explosions and ends with the formation of an extrusive lava dome and the emission of lava flows (Melekestsev et al., 1995). These cycles differ in the volumes and compositions of erupted rocks, in eruptive parameters, in magmatic productivity, and in the ratio of juvenile to resurgent material and are dated, respectively, between 12,000-11,000 to 8400 years ago, 8400-6400, 6400-3000, and 3000 years ago to the present (Melekestsev et al., 1995). Kizimen is known to be a source of heterogeneous, poorly mixed magmas with compositions ranging from dacites to basaltic andesites (Melekestsev et al., 1995;[START_REF] Piip | Kizimen volcano[END_REF]. The morphology of Kizimen is a cone with lava domes and thick lava flows surrounded by volcaniclastic material (Melekestsev et al., 1995) and it is similar to Mount St. Helens volcano before its 1980 eruption [START_REF] Lipman | The 1980 eruptions of Mount St[END_REF][START_REF] Shakirova | Features of the Kizimen Volcano area seismicity prior to and during the 2010-2013 eruption[END_REF]

(Figure II.19).
The four cycles of activity will be detailed hereafter, based on the study of Melekestsev et al. (1995).

The activity of the KZI cycle reached its peak of activity about 10,000 years ago during a series of explosive eruptions with an estimated volume of 1 km 3 . KZII (8400 to 6400 BP) began with a series of moderate explosive eruptions, with volumes of 1103 to 1102 km 3 . A large Plinian eruption then occurred, dated first at 8,300 years by Melekestsev et al. (1995) and re-estimated at 7500 to 7600 years by Braitseva et al. (1992[START_REF] Braitseva | Ages of calderas, large explosive craters and active volcanoes in the Kuril-Kamchatka region, Russia[END_REF]. This eruption involved 2.5 to 3 km 3 of magma and deposits were observed several hundred of km from Kizimen (as on Bering Island). This eruption is considered as one of the most important in the history of Kamchatka and its deposits are used as marker in tephrochronology on the scale of the whole Kamchatka peninsula (Melekestsev et al., 1995). The KZII cycle ended with the growth of a second lava dome, followed by a 1200-year period of inactivity. During the third eruptive cycle of Kizimen (KZIII, 6400-3000 years ago), the eruptive activity of Kizimen is very modest compared to the previous two cycles (eruptive intensity of one order of magnitude lower than in cycle KZII). The initial eruption of the cycle is the most important with about 0.1 km 3 of resurgent and juvenile material emitted, with coarse blocks. This eruption is followed by a series of eruptions that emitted 1.1-2 km 3 of ash and tephra fallout and the cycle ends with a period of about 3000 years of repose time. The last cycle of activity of Kizimen (KZIV) began 3000 years ago and continues until present, with the activity of the volcano being closer to what it was during the KZI and KZII cycles. It was initiated by an important explosive eruption, emitting pyroclastic density currents and growth of a large extrusive lava domes with lava flows. The magma compositions of the emitted magmas were more basic than compared to the whole history of the volcano (Melekestsev et al., 1995). This eruption was followed by a series of other eruptions whose volume reached 0.7 to 1 km 3 . A series of lava domes were set up over a period of approximately 1100 years, in parallel with an explosive and effusive activity. The activity of Kizimen at that stage was similar to that of modern Bezymianny (Melekestsev et al., 1995).

The volcano produced an eruption in 1927-1928 that was probably of hydrothermal origin and where no deposits were found (Churikova et al., 2007). In 1963, the volcano was the site of intense seismic volcano-tectonic earthquakes activity (earthquake swarm of Shchapinsky; earthquakes from the ground surface up to 35 km depth) (Senyukov et al., 2011) (Figure II.19;Figure II.20). 130 earthquakes were catalogued during this period, with two major ones on 26 and 27 May 1963 (20-25 km to the west from Kizimen [START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF]. They were followed by an increase of the fumarole activity [START_REF] Petrov | Modern sulphate activity of Kizimen volcano[END_REF]. Isolated events were then observed for several decades [START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF]. In 2010-2013, Kizimen produced its first historical magmatic eruption (0.4 km 3 DRE of magma emitted) with explosive episodes generating pyroclastic density currents (0.27 km 3 ) followed by the extrusion of a lava dome that gradually formed a very thick lava flow (0. Melekestsev et al. (1995) noticed a regular increase in more basic compositions towards younger eruptions at Kizimen but it ended after a long dormancy with the 2010-2013 eruption, which produced the most evolved compositions in the record of Kizimen volcano (the previous magmatic eruption of the volcano erupted lavas of basaltic andesite composition with 54.5% SiO2) (Auer et al., 2018). II.3.5.1. 2010II.3.5.1. -2013 eruption eruption A seismic crisis with more than 3000 seismic events were detected from 01.04.2009-01.05.2011 and in October 2010, three strong earthquakes occurred at 2 km above sea level and up to 3 km depth (magnitude of 5) [START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF] Therefore, it can be assumed that unless there are new strong earthquakes of magnitude > 4 in the present Kizimen activity, subsequent Kizimen eruptions in this current phase will be no stronger than those that have already occurred [START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF].

This eruption will be studied in this thesis. A model proposed by Auer et al. (2018) for the 2010 eruption suggests the presence of a silicic magma chamber located at 5-11 km depth (Figure II.22). The dacitic magma is estimated to have pre-eruptive temperatures of ~810 °C, fO2 of 0.9-1.6 log units above the NNO buffer (5.5 wt% H2O) and shallow crustal storage conditions at ~123 MPa (~5 km) (Auer et al., 2018). They propose that an injection of basaltic magma into an evolved dacitic reservoir occurred at depth in 1963, associated with a seismic crisis, which was previously detailed (Figure II This seismic swarm is thought to have been caused by the injection of a dike approximately 15 km long and 20 cm thick beneath the volcano according to interferometric synthetic aperture radar (InSAR) ground deformation data (Auer et al., 2018;Ji et al., 2013). This dike would have propagated vertically from a depth of about 10 km (Ji et al., 2013), but as no deep seismicity was recorded, Auer et al. (2018) think that it is difficult that such intrusion could occur aseismically. Indeed, for the neighboring Tolbachik volcano, intrusion of dikes generate multiple earthquakes at depth up to 30 km (Auer et al., 2018). An increase in the temperature of the Kizimen fumarole was also observed (from 270 to 340 °C) by Tembrel and Ovsyannikov (2009). But these authors do not believe that this is due to seismic reactivation (Tembrel and Ovsyannikov, 2009). Rapid mixing of magma components in a probably stratified reservoir is thought to have caused the eruption in November 2010, ejecting poorly mixed magma evidenced by abundant andesites and banded dacites (Auer et al., 2018) (Figure II.22). The last mixing event happened alongside the start of the eruption, suggesting that the eruption itself induced mixing, which is an observation that had previously been made on other volcanoes such as Shishaldin volcano in Alaska or Etna in Italy (Auer et al., 2018;Costa et al., 2020)

(Figure II.22).

There were few highly explosive events involving the eruption of highly vesiculated magma, in the Late Holocene history of Kizimen (Melekestsev et al., 1995). The scenario of the 2010-2013 eruption can be considered as typical for Kizimen as the most common one represents slow extrusion of highly viscous, partly degassed silica-rich magma accompanied by only moderate explosive activity (Auer et al., 2018). One hypothesis made by these authors to explain the fact that Plinian explosive stages are rather weak at Kizimen, compared to other volcanoes with similar compositions and whose eruption would start with a highly Plinian stage, is that the silicic magma chamber is partly degassed long before eruption, with magmatic fluids leaking through the active tectonic fault zone that crosses the edifice.

The fault zone probably favors the continuous leakage of volatiles from the shallow dacitic magma chamber, which feeds the fumarole on the flank of the volcano and prevents it from reaching an eruptible state (Auer et al., 2018). This passive degassing can depressurize the magma chamber and increase the viscosity of the magmas in the shallow chamber, which lowers the explosivity of its eruptions (Auer et al., 2018;Girona et al., 2015Girona et al., , 2014)). Previous studies showed the high viscosity of the magmas (Auer et al., 2018;Churikova et al., 2007;Melekestsev et al., 1995;[START_REF] Piip | Kizimen volcano[END_REF]. This high viscosity can also explain the poorly intermingled magmas, acting as a density filter for mafic intrusions from depth and can explain the long period between the injection of magma at depth and the eruption (model of Auer et al. (2018)). Shallow earthquake swarms before Kizimen eruptions can then last more than a year due to the high viscosity of the ascending magma (Auer et al., 2018). A future gravitational failure of the volcanic edifice is very likely due to the high viscosity of the magmas and Kizimen location on an active tectonic fault. The character of explosive activity will depend on the depth of magma inside the edifice at the moment of failure. It has been previously noted by Melekestsev et al. (1995) for Kizimen, whom suggest that the future eruption style may be similar to the lateral blast of 1980 of Mount St. Helens (Auer et al., 2018;[START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF]Melekestsev et al., 1995). II.3.5.3. Monitoring In April 1963, earthquakes near Kizimen volcano began to be recorded, there were nine stations in Kamchatka at that time, the nearest station from Kizimen was "Kozyrevsk" (~100 km away) [START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF] [START_REF] Shakirova | Features of the Kizimen Volcano area seismicity prior to and during the 2010-2013 eruption[END_REF]. However, there is no ground-based deformation monitoring network on Kizimen (too remote, inclement weather and difficult logistics) but remote sensing techniques are also applied on Kizimen (InSAR data analysis) (Ji et al., 2013). In July 2011, the Kamchatka Branch of the Russian Geophysical Survey installed an automatic monitoring camera on TUMD with the frame rate of one image per minute near Tumrok hot springs located 10 km to the NNE from the volcano, to have complete visual data on the eruptive processes (Auer et al., 2018;Firstov and Shakirova, 2014).

CHAPTER III: SCIENTIFIC APPROACH, SAMPLES AND ANALYTICAL METHODS

Résumé

Dans ce troisième chapitre, l'approche scientifique et les principales méthodes utilisées durant ma thèse sont présentées. Dans un premier temps, les échantillons et la démarche d'échantillonnage sont abordés. Les échantillons de l'île de la Dominique proviennent de missions sur le terrain qui ont eu lieu avant ma thèse, en 2011 et 2013 et comprennent des échantillons de retombées de ponces correspondant à des éruptions explosives ponceuses du volcan Morne Trois Pitons-Micotrin (Boudon et al., 2017). Les éruptions étudiées sont celles de Grand Fond (24 ka cal BP), une éruption ayant mis en jeu un grand volume de magma (éruption ignimbritique) et quatre éruptions Pliniennes ayant mis en jeu de plus petits volumes de magmas (Goodwill Entrance : 18 ka cal BP, "Plinian Post Roseau" 1 (PPR1) : 12 ka cal BP, PPR2 : 12 ka cal BP et PPR3 : 9 ka cal BP). L'échantillonnage au Kamchatka s'est déroulé en 2019 lors d'une campagne de terrain sur les volcans du Bezymianny (phase paroxysmale de l'éruption de 1956, principalement) et du Kizimen (éruption de 2010-2013). Les échantillons correspondent pour le Bezymianny : i) à des fragments du cryptodôme décapité lors de l'explosion latéralement dirigée de 1956 suite à la déstabilisation du flanc est du volcan, ii) à des retombées de ponces de la phase Plinienne ayant suivi l'explosion latéralement dirigée. Pour le Kizimen, les échantillons correspondent à des fragments prélevés dans les dépôts de courants de densité pyroclastiques provenant de phases d'écroulement du dôme de lave et du front de la coulée de lave épaisse ayant accompagné l'éruption extrusive/explosive de 2010-2013. Ces échantillons sont représentatifs de l'éruption et permettent d'établir les conditions de stockage et les processus magmatiques pré-éruptifs auxquels nous nous intéressons dans cette thèse.

dans les magnétites est aussi abordée, moins en détails que pour les orthopyroxènes, ainsi que le modèle de diffusion du Ti.

Par la suite, la méthode de détermination de l'architecture et de la dynamique spatio-temporelle du système magmatique est présentée : la « Crystal System Analysis » (CSA), pour l'identification d'environnements magmatiques et des processus mis en jeu en profondeur, qui, couplée à la diffusion intracristalline dans les orthopyroxènes, nous amène à déterminer les échelles de temps des processus magmatiques pré-éruptifs (Kahl et al., 2011). Ce type d'étude nécessite de traiter un nombre suffisant de cristaux afin d'en tirer des informations représentatives du système étudié.

-Bezymianny: 5 samples from the 1956 climactic phase, a benchmark eruption for flank collapse, related to cryptodome emplacement and then followed by a laterally directed blast, were collected (see Chapter II. 3.4;Figure III.2a). Fragments of the andesitic cryptodome, of different vesicularities, were collected from several sites on the eastern flank of the volcano in the directed blast deposits (Table III.1). These blast deposits ("blast surge") are preserved in various areas in the proximal zone (less than 19 km from the volcano) and show a three-layer stratification (A, B, and C, Figure III.3), characteristic of the dilute, turbulent pyroclastic density current [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF]) (Figure III.4a). Debris avalanche from the flank destabilization preceded the blast but as the blast surge had a speed of more than 60 m/s at 10 km from the volcano, the surge overran the debris avalanche at this distance [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF]) (Figure III.2). Additional pumiceous fragments from the Plinian phase that followed the 1956 flank collapse were also collected from more distal parts of the southeastern flank of the volcano (Figure

III.4b).

At low points and in channels, most of the deposits are either covered by other eruptions or eroded. between 5 and 10 cm, they are reduced in fragments of 3 cm for a finer crushing. They are first cleaned in an ultrasonic tank to remove any adherent matrix, then dried at 80 °C for ~48 h. Each crushed sample is quartered. One part is used for a finer grinding in order to perform whole rock analysis (see III.

2.4.1).

The other part is sieved with a 2 mm to 125 µm sieve column (2-1 mm, 1000-710 µm, 710-500 µm, 500-355 µm, 355-250 µm, 250-125 µm or 500-315 µm, 315-250 µm) for crystal selection. The different fractions are then ultrasonically cleaned with ultrapure water (Milli-Q®) and dried at 80 °C for ~48 h. They are then observed under a binocular magnifier in order to carry out specific investigations on crystals and to determine in which fraction(s) the studied crystals are the most abundant and the most automorphic.

These crystals are selected by hand in different fractions, according to the type of crystals and their size (for example, orthopyroxenes and iron oxides crystals generally appear the most automorphic and are well preserved in the 500-355 and 355-250 µm fractions), then, they are aligned on a tape on a thin section, with orthopyroxenes having their c-axis in north-south direction. The crystals are surrounded by a plastic ring and then impregnated in Epoxy resin (at room temperature and under ambient pressure conditions). They are then dried for at least 9 hours. These mounts are then polished from 22 to 0.3 µm so that the crystals contained in the mounts are cut in their middle (crystals of similar size and thickness) and that their surface is perfectly smooth and flat. In order to check the quality of the polishing and the surface of the mount, regular checks are made under a binocular magnifier and an optical microscope The electron microprobe is based on the same principle as the SEM but it allows to acquire quantitative chemical analysis of crystalline phases and glasses by analyzing the X-rays emitted by the samples under their surface following their excitation by an electron beam. X-rays are produced when incident electrons pass through the sample, creating secondary electrons, leaving "holes" in the inner electron layers of atoms. To stabilize these atoms, electrons from the outer electron layers will jump to the inner layers, however, since the outer layers are of higher energy, the atom loses energy in the form of X-rays (Figure III.8). These X-rays have energies that are specific to the elements contained in the samples (Goldstein et al., 2003). Since the portion of the sample interacting with the incident electron beam corresponds to a specific volume, special attention must be paid to the step between the analysed points of an electron microprobe transect in order to avoid convolution effects between the points.

III.2.4.3. Analytical conditions and uncertainties

We used electron microprobes (CAMECA SX-Five and SX-100) from the Service Camparis (Paris) to measure the major and minor element compositions (Si, Ti, Al, Fe, Mg, Mn, Ca, Na, K, P and Cl) of the crystals (orthopyroxenes, plagioclases, Fe-Ti oxides, amphiboles), melt inclusions and residual glasses. Before the analyses, the electron microprobe was calibrated with internal standards whose compositions are known.

The analytical conditions are an accelerating voltage of 15 kV, a beam current of 10 nA and a focused beam (size < 2 µm) for the crystals. The counting times of some elements (Fe and Mg) that are of direct interest for the study of zonations in orthopyroxenes have been increased to 80 s in order to have a better accuracy (time of 60 s for Fe, Ti and Al for the profiles in magnetites and 30 s for Fe, Ti and Al for the points in the magnetites/ilmenites couples). The composition profiles on zoned crystals have a 2 µm step. Unzoned crystals were analyzed on a spot basis (4 points per crystal).

For the measurement of major and minor elements in the residual glasses of the samples, an acceleration voltage of 15 kV was used, with counting times varying from 10 to 25 s depending on the element and a beam size of 6-10 µm to avoid alkaline migration (Balcone-Boissard et al., 2018;[START_REF]Evidence for an Active, Transcrustal Magma System in the Last 60 ka and Eruptive Degassing Budget (H 2 O, CO 2[END_REF]Solaro et al., 2019). To avoid Na loss, a beam current of 4 nA was used and Na and Si elements were counted first and simultaneously. The internal standards used for intercalibration of glass measurements by EPMA were analyzed under the same conditions as the residual glasses (Corbetti volcano obsidian, Ethiopia (CO5), Lipari obsidian, Aeolian Island, Italy (LIP) and Little Glass Mountain, California, USA (LGM) [START_REF] Balcone-Boissard | Non-volatile vs volatile behaviours of halogens during the AD 79 plinian eruption of Mt. Vesuvius, Italy[END_REF].

Inclusions <10 µm could not be analyzed, nor could fractured zones in zoned crystals.

Analytical errors on major and minor elements were calculated to be 0.4 wt% for Si, 0.07 wt% for Ti, Al, Ca, Na, K, and Ni, 0.2 wt% for Fe, Mg, and Mn, and 0.04 wt% for Cr and P.

III.3. Diffusion chronometry

When a crystal grows in the magma reservoir, it is surrounded by a liquid with specific conditions of temperature, pressure, oxygen fugacity and arrival of volatiles [START_REF] Kahl | Compositionally zoned crystals and real-time degassing data reveal changes in magma transfer dynamics during the 2006 summit eruptive episodes of Mt. Etna[END_REF], 2011) (Figure III.10a). These stable storage conditions define a given magmatic environment (ME) and are recorded during crystal growth. However, upon changes occurring in the magma chamber (from magma mixing to magma degassing, magma transport, convection or fractionation), these storage conditions change and can be recorded as chemical zonations in the crystals (Figure III.10b) [START_REF] Anderson | Probable relations between plagioclase zoning and magma dynamics, Fuego Volcano, Guatemala[END_REF][START_REF] Giacomoni | Plagioclase as archive of magma ascent dynamics on "open conduit" volcanoes: The 2001-2006 eruptive period at[END_REF]Petrone and Mangler, 2021;[START_REF] Streck | Mineral Textures and Zoning as Evidence for Open System Processes[END_REF]. These zonations or borders of different compositions are evidence of changes in pre-eruptive conditions in the storage system but they can also be formed as a result of crystallization (leading to melt differentiation and the development of crystal zonation) (Costa et al. 2020). But the chemical gradients that make the zonations in the crystals are unstable and tend to fade due to kinetic diffusion processes by chemical element diffusion [START_REF] Blundy | Trace element evidence for plagioclase recycling in calc-alkaline magmas[END_REF]Costa et al., 2008;Costa and Morgan, 2010;[START_REF] Ginibre | Crystal Zoning as an Archive for Magma Evolution[END_REF]Ginibre et al., , 2004Ginibre et al., , 2002aGinibre et al., , 2002b;;[START_REF] Humphreys | Magma Evolution and Open-System Processes at Shiveluch Volcano: Insights from Phenocryst Zoning[END_REF][START_REF] Kahl | Compositionally zoned crystals and real-time degassing data reveal changes in magma transfer dynamics during the 2006 summit eruptive episodes of Mt. Etna[END_REF]Kahl et al., , 2011;;Petrone and Mangler, 2021;Ruprecht and Wörner, 2007;Tomiya and Takahashi, 2005;[START_REF] Wark | Pre-eruption recharge of the Bishop magma system[END_REF]

(Figure III.10c).
The extent to which diffusion tends to balance the chemical zonation depends on the mineral, the elements under consideration, temperature, and time (Costa and Morgan, 2010). If the diffusion coefficients (at magmatic temperatures) of the elements of interest are known, diffusion re-equilibration times can be calculated using Fick's second law or other more complex models (Costa and Morgan, 2010;[START_REF] Streck | Mineral Textures and Zoning as Evidence for Open System Processes[END_REF].

Crystals have long been used as records of magmatic processes and the timescales found by diffusion modelling are in the order of hours, days, years to hundreds of years that can temporally constrain events such as magma injection prior to an eruption, whereas with radiochronology, these are on the order of thousands of years (Costa et al., 2020;Costa and Morgan, 2010).

Since the 1950s, with the development of chemical analyses such as SEM [START_REF] Castaing | Application des sondes électroniques à une méthode d'analyse ponctuelle, chimique et cristallographique[END_REF], or EPMA as well as other techniques more recently, such as laser ablation mass spectrometry (LA-ICP-MS), chemical zoning is also studied as gradients of trace elements and isotopic ratios (Costa and Morgan, 2010;[START_REF] Davidson | Microsampling and Isotopic Analysis of Igneous Rocks: Implications for the Study of Magmatic Systems[END_REF][START_REF] Ginibre | Crystal Zoning as an Archive for Magma Evolution[END_REF]. A chemical gradient encourages mixing in any miscible system, with the lowering of the free energy of the system through a mixing process. Interactions can appear between different ions as various crystallographic sites can be occupied by different elements with different charges which produces a multi-component diffusion process with a diffusion coefficient characteristic for each component [START_REF] Lasaga | Multicomponent exchange and diffusion in silicates[END_REF].

III.3.1.1. Diffusion equations

To obtain a net flux or directed motion due to diffusion, a driving force that pushes particles, such as atoms, molecules or ions, is required to obtain motion in a preferred direction (Costa et al., 2008). In the case of matter transport, the presence of a driving force in the medium, such as a chemical potential gradient, causes matter to flow in a specific direction. The direction of the flux J is in the opposite direction of the chemical potential gradient. The chemical potential is the change in energy between states following an act of mixing, for example (Costa and Morgan, 2010). These laws, in the form of partial differential equations show that the flux of a component i depends on the chemical potential gradient 𝜕𝜇 𝑖 𝜕𝑥 (for one dimension, via a proportionality constant Li) and the type of material, following equation ( III.3):

𝐽 𝑖 = -𝐿 𝑖 𝜕𝜇 𝑖 𝜕𝑥 (III. 3)
However, since chemical potential is not easily observable or measurable, a relationship between chemical potential and measurable concentrations can be defined by defining a reference state and a reference chemical potential 𝜇 0 , following equation (III.4):

𝜇 𝑖 = 𝜇 0 + 𝑅𝑇 𝑙𝑛(𝛾 𝑖 𝐶 𝑖 ) (III. 4)
By combining (III. 4) and (III.8), we obtain:

𝜕𝐶(𝑥, 𝑡) 𝜕𝑡 = 𝜕𝐶 𝜕𝑥 (𝐷 𝜕𝐶(𝑥, 𝑡) 𝜕𝑥 ) (III. 10)
If the diffusion coefficient D is constant with distance x, the previous equation can be rewritten: . 11) with D, the diffusion coefficient independent of the position or the concentration C.

𝜕𝐶(𝑥, 𝑡) 𝜕𝑡 = 𝐷 𝜕 2 𝐶(𝑥, 𝑡) 𝜕𝑥 2 (III
A frequent scenario in nature is the competition between the phenomena of diffusion and growth/dissolution of crystals. In the case of a crystal growing or dissolving in a liquid, the total flux is the result of the diffusion and the flux coming from the growth or dissolution of the crystal in the medium (liquid for example), at the speed v. This type of diffusion is therefore accompanied by the phenomena of growth or dissolution and is called the "moving boundary" problem (Costa et al., 2008).

Applications of diffusion with growth/dissolution models are quite limited (Costa et al., 2008;Nakamura, 1995b;[START_REF] Ozawa | Kinetics of diffusion-controlled evaporation of Fe-Mg olivine: experimental study and implication for stability of Fe-rich olivine in the solar nebula[END_REF][START_REF] Watson | Dissolution, growth and survival of zircons during crustal fusion: kinetic principles, geological models and implications for isotopic inheritance, in: The Third Hutton Symposium on the Origin of Granites and Related Rocks[END_REF].

III.3.1.2. Solutions of the diffusion equation

Fick's second law describes the changes in concentration with time at each position for a concentration profile. It can be used to obtain the time that a crystal spends at a given temperature, if gradients of chemical compositions are present in a crystal or liquid. However, the following points must be considered. First, the way of characterizing the concentration gradients and the spatial resolution must be considered, but also the values of the diffusion coefficient, as well as the initial conditions (initial shape of the concentration profile) and the boundary conditions (Costa and Morgan, 2010). To obtain the concentration 𝐶(𝑥) at each time, after the onset of diffusion, it is necessary to add three equations:

-an initial condition from which the diffusion process will start, corresponding to the shape of the composition profiles before the diffusion starts (four types) (Costa et al., 2020) (Figure III.13).

-two other equations providing constraints on the boundary concentrations of the system (boundary conditions) (Costa et al., 2008). Boundary conditions derive from the nature of the exchanges of elements of interest at the edge of the crystal with the surrounding matrix. They control the exchange of matter between the crystal and its environment (open boundaries) or not (closed boundaries) (Costa et al., 2020) In our case, to study the diffusion process in a crystal, a boundary condition is necessary to specify the concentration or flux evolution at the crystal surface. If this crystal is in equilibrium with a liquid at a constant temperature, the concentration at the surface of the crystal will remain constant and remain at equilibrium, while diffusion will tend to homogenize the concentration gradients in the crystal (Costa et al., 2008)

(Figure III.10).
The equation ( 9) can be solved in different ways, on the one hand with numerical methods (like the finite difference method) or analytically, for example with the error function Erfc for a plane solution and in an infinite medium (Costa et al., 2020(Costa et al., , 2008;;Costa and Morgan, 2010;[START_REF] Crank | The Mathematics of Diffusion[END_REF]. If the initial and boundary conditions and the diffusion coefficient of the element of interest are known, the time can be obtained by fitting the model to the concentration profile [START_REF] Browne | The influence of magma ascent path on the texture, mineralogy, and formation of hornblende reaction rims[END_REF]).

In the most common geological contexts, three analytical solutions can be used: a planar source with a thin surface layer (in a finite solid) [START_REF] Philibert | Diffusion et transport de matière dans les solides[END_REF], a constant superficial concentration (diffusion in a semi-infinite medium) (Figure III.15a) or an infinite initial distribution (Figure III.15b) [START_REF] Philibert | Diffusion et transport de matière dans les solides[END_REF]Solaro-Müller, 2017). This last analytical solution is used to model the interdiffusion of Fe-Mg elements in orthopyroxene crystals for my thesis. From a step profile, the evolution of the compositional profile as a sigmoid is modeled as a function of time.

For a planar source, with an infinite initial distribution, two media of different concentrations C are brought into contact (initial conditions: C1 for x > 0 and C2 for x < 0). These conditions define a staircase profile, symmetrical and with a concentration at the interface (for x = 0), equal to

1 2 (C 1 + C 2 ).
The solution of the diffusion equation can be written as:

𝐶 -C 1 C 2 -C 1 = 1 2 𝐸𝑟𝑓( 𝑥 2√𝐷𝑡 ) (III.
12)

The diffusion pattern can be determined for each time interval, using, for example, the least square root method and the mismatch between each time step can be used to determine the best time model [START_REF] Girona | DIPRA: A user-friendly program to model multi-element diffusion in olivine with applications to timescales of magmatic processes[END_REF].

III.3.1.3.b. Pressure

Diffusion processes in solids are pressure dependent, due to the pressure dependence of the Gibbs free energy of an atom to make a diffusion jump. Pressure influences the activation volume of diffusion ∆𝑉, which is the change of volume in the crystal following a jump of an atom from one site to another at constant temperature [START_REF] Mehrer | The Effect of Pressure on Diffusion[END_REF]. Pressure tends to increase order in geological systems by decreasing the mobility of cations and changing the diffusion mechanism of cations in the crystal lattice [START_REF] Hazen | Effects of pressure on order-disorder reactions[END_REF]. The effect of pressure on diffusion can be described according to the following relationship:

𝐷 0 𝛼 (𝑒 𝑃𝑉 𝑅𝑇 ) (III. 15)

III.3.1.3.c. Oxygen fugacity

The oxygen fugacity fO2 is of the dimension of a pressure, it is the equivalent of the partial pressure of oxygen in a particular environment (rocks, atmosphere...) corrected for the non-ideal character of the gas [START_REF] Albarède | Oxygen Fugacity[END_REF]. The amount of oxygen available in a system can influence the mineralogical structure of a system [START_REF] Bonin | Magmatisme et roches magmatiques[END_REF]. For a given redox couple, an increase in fO2 leads to a shift of the equilibrium towards more oxidized species [START_REF] Pinet | Les phénomènes d'oxydoréduction dans les verres d'oxyde[END_REF]. There are mineral assemblages, redox couples with a buffering effect (stabilizing effect), where the oxygen fugacity depends on the temperature. Thus, in experimental petrology, the oxygen fugacity is measured during isobaric processes to better explain the stability of minerals [START_REF] Bonin | Magmatisme et roches magmatiques[END_REF].

The main redox couples with buffering capacity (experimental studies carried out at constant pressure of 1 bar) are ranked here in descending order of oxygen fugacity [START_REF] Zhao | An oxygen barometer for rutile-ilmenite assemblages: oxidation state of metasomatic agents in the mantle[END_REF] :

- The oxygen fugacity parameter is important to consider as the crystal lattice can change depending on the oxidation state of the system. For iron-containing minerals, such as in olivines or orthopyroxenes, an increase in oxygen fugacity increases the amount of octahedral vacancies due to the partial oxidation of Fe 2+ in Fe 3+ which is balanced in terms of charge by the formation of these vacancies V: 3 Fe 2+ → 2 Fe 3+ + V (Dohmen et al., 2007). This increases the diffusion volume (Solaro-Müller, 2017). For Fe-Mg rich silicate minerals, the exact dependence between the diffusion coefficient and oxygen fugacity is expressed as an exponential (Allan et al., 2013;[START_REF] Buening | Fe-Mg lattice diffusion in olivine[END_REF][START_REF] Chakraborty | Compositional Zoning and Cation Diffusion in Garnets[END_REF]Dohmen et al., 2016Dohmen et al., , 2007;;Ganguly and Tazzoli, 1994):

𝐷 𝛼 (𝑓𝑂 2 ) 𝑚 (III. 16)

III.3.1.3.d. Water fugacity

The effect of water or any other fluid can influence diffusion in a medium if it affects the properties of that medium. Since the vacancies in silicates are charged (for example, a missing Mg 2+ ion will result in a doubly negatively charged vacancy), the incorporation of water or other fluids will therefore balance or unbalance the structure, thus affecting the diffusion rates by changing the amount of vacancies present (Costa et al., 2008).

The effect of water on diffusion rates can be evaluated if the amount of water incorporated depends on the water fugacity. For olivine, if sufficient concentrations of hydrogen are incorporated into the crystal, the charge balance may change (vacancies balanced by hydrogen atoms rather than by Fe 3+ ). This would lead to a change in the diffusion, going macroscopically from a "dry" to a "wet" state (Costa et al., 2008;Costa and Chakraborty, 2008).

The diffusion coefficient in a water bearing system is expressed as:

𝐷 𝑊 = 𝐴 𝑊 𝑓𝑂 2 𝑛𝑤 (𝑓𝐻 2 𝑂) 𝑟 𝑎 𝑆𝑖𝑂 2 𝑚𝑤 𝑒 (-𝐸 1𝑤 +(𝑃-10 5 )∆𝑉 𝑊 𝑅𝑇 ) (III. 17)
With w indicating the "wet" state of the system (Costa and Chakraborty, 2008).

III.3.1.3.e. Mineral compositions and anisotropy

The determination of the diffusion coefficient can also depend on the composition of the crystals and their possible anisotropy. The effect of anisotropy depends on the inner structure of the minerals and thus on the jump required for the elements in the different directions. The diffusion will be faster in the direction for which the jump is the smallest, for an anisotropic crystal, thus, depending on the crystallographic direction, two or three diffusion coefficients can be calculated for orthorhombic, hexagonal, trigonal or tetragonal crystals [START_REF] Brady | Diffusion in Minerals: An Overview of Published Experimental Diffusion Data[END_REF]Solaro-Müller, 2017).

Thus, for some silicates such as plagioclases, pyroxenes or olivines, the enthalpy of formation of the vacancies within the crystal structure and thus the diffusion coefficient of certain elements, depends on the composition of the crystal. For plagioclases, the diffusion coefficients of the elements Mg, Sr, Ba and Pb are greater for plagioclases rich in Na (albite) whereas they are smaller for plagioclases rich in Ca (anorthite) [START_REF] Blundy | Crystal-chemical controls on the partitioning of Sr and Ba between plagioclase feldspar, silicate melts, and hydrothermal solutions[END_REF][START_REF] Van Orman | Magnesium diffusion in plagioclase: Dependence on composition, and implications for thermal resetting of the 26 Al-26 Mg early solar system chronometer[END_REF]. Ferromagnesian crystals such as orthopyroxenes and olivines show a dependence of the diffusion on the molar fraction of Fe: XFe. Indeed, as the element Fe can be present in these two oxidized states, it favors charge balancing when a vacancy forms; the crystals having a composition richer in Fe have therefore more vacancies in their structure (Solaro-Müller, 2017).

The effect of composition on diffusion can be written as:

𝐷 0 𝛼 𝑒 𝛼𝑋 𝑖 (III. 18)
Where 𝛼 is constant and 𝑋 𝑖 is the molar fraction of the element i constituting the mineral.

III.3.1.3.f. Final diffusing equation

The final form of the diffusion coefficient equation is therefore:

𝐷 = 𝐷 0 (𝑒 αXi+ -𝐸 𝑎 +𝑃𝑉 𝑅𝑇 ) (𝑓𝑂 2 ) 𝑚 (𝑓𝐻 2 𝑂) 𝑛 (III. 19)
Where 𝐷 0 is the pre-exponential factor (m 2 .s -1 ), Q or 𝐸 𝑎 , the activation energy at 10 5 Pa (J.mol -1 ), α constant, V the activation volume (m 3 .mol -1 ), T the temperature (K), P pressure (Pa), R the ideal gas constant (J.mol -1 .K -1 ) and fO2 and fH2O, the oxygen and water fugacity, respectively. As seen before, the diffusion can be faster along a specific crystallographic axis, so several diffusion coefficients can be calculated, depending on the crystal structure [START_REF] Brady | Diffusion in Minerals: An Overview of Published Experimental Diffusion Data[END_REF]Solaro-Müller, 2017).

III.3.2.1. Diffusion coefficient estimation in the pyroxenes

Direct and indirect measurements of diffusion in pyroxenes are based on the one hand on chemical or isotopic analysis of profiles as a function of depth in the crystal using an electron microprobe (coupled to Energy Dispersive X-ray (EDX)), accelerator-based ion beam techniques or ion probe (SIMS) analyses (Cherniak and Dimanov, 2010). On the other hand, indirect estimates involve experimental measurement of the kinetics of lamellar exsolution homogenization, order-disorder at crystal sites and solid-state reactions with moving interfaces (Cherniak and Dimanov, 2010).

The first studies on the interdiffusion of major divalent cations (Fe, Mg, Ca) in pyroxenes were carried out in order to constrain the rate of cooling and the duration of "annealing" (thermal process modifying the physical and chemical conditions of a material) of extraterrestrial assemblages, studying lunar basalts (Cherniak and Dimanov, 2010;[START_REF] Huebner | Diffusively rimmed xenocrysts in 77115[END_REF][START_REF] Huebner | Cooling history and significance of exsolved pyroxene in lunar noritic breccia 77215[END_REF][START_REF] Huebner | Assessment of Diffusion in Pyroxenes: what we do and do not Know[END_REF][START_REF] Stanford | Reexamination of diffusion processes in 77115 and 77215[END_REF]. But it is only since the 1980s and the development of microscopic and deep analytical techniques (SIMS, Rutherford Backscattering Spectroscopy (RBS), NRA) and their high spatial resolution (~15-30 nm) that short diffusivities could be investigated (possibility to measure short diffusion profiles) (Cherniak and Dimanov, 2010). Some studies even consider the effect of sample composition, pressure, crystal structure and its possible anisotropy (Cherniak and Dimanov, 2010).

Concerning Fe-Mg interdiffusion in orthopyroxenes, the first estimate of these diffusion rates was made by Ganguly and Tazzoli (1994). Their study is based on the experimental determination of the distribution kinetics of Fe and Mg elements between the M1 and M2 sites of the crystal structure [Fe 2+ (M2) + Mg (M1) ⇔ Fe 2+ (M1) + Mg (M2)] [START_REF] Anovitz | Order-disorder experiments in orthopyroxenes: Implications for the orthopyroxene geospeedomete[END_REF][START_REF] Besancon | Rate of cation ordering in orthopyroxene[END_REF]Ganguly and Tazzoli, 1994;[START_REF] Saxena | Kinetics of Fe 2+ -Mg distribution in aluminous orthopyroxenes[END_REF]. Based on crystal structure data of orthopyroxenes, the diffusion in orthopyroxenes is considered to be anisotropic in the Ganguly and Tazzoli (1994) interpreted the rate constant for disordering and ordering processes between the two sites as an indicator of the diffusion coefficient in terms of Fe-Mg crystal interdiffusion, thus establishing an equivalence between the interdiffusion constant and the order-disorder constant (Cherniak and Dimanov, 2010).

The parameterization of Ganguly and Tazzoli (1994) of the Fe-Mg interdiffusion coefficient in orthopyroxenes along c-and b-directions of orthopyroxene is as follows:

Log D (𝐹𝑒-𝑀𝑔)𝑏-𝑐 = -5,54 + 2.6𝑋 𝐹𝑒 - 12530 T (III. 20)
Where 𝑋 𝐹𝑒 is the composition of the crystal in terms of molar percentage of iron and T the temperature in Kelvin. The experiments are performed for an iron composition of the enstatite crystals between 0.1 and 0.5 XFe, a temperature between 500 and 800 °C, for some experiments the oxygen fugacity is controlled by a mixture of H2 and CO2 between 0.5 and 0.8 log units above the IW buffer, while for others, these conditions were defined by vacuum conditions in sealed capsules (Cherniak and Dimanov, 2010;Ganguly and Tazzoli, 1994).

Moreover, the presence of oxygen in pyroxenes have effects on iron by oxidizing it (Fe 2+ and Fe 3+ ), so the oxygen fugacity has an influence on the diffusion. Knowing that the equation of the diffusion coefficient is of this form (see parts above):

D = 𝐷 0 exp (α𝑋 𝑖 + -𝐸 𝑎 + 𝑃𝑉 𝑅𝑇 ) 𝑓𝑂 2 𝑚 𝑓𝐻𝑂 2 𝑛 (III. 21)
By analogy with studies in olivines [START_REF] Buening | Fe-Mg lattice diffusion in olivine[END_REF], Ganguly and Tazzoli (1994) conjectured that the oxygen fugacity dependence was of the same order, with exponent m ~1/6 (same mechanism for generating vacancies in the crystal structure upon oxidation by iron). Stimpfl et al. (2005) studied the effect of oxygen fugacity on order-disorder kinetics and estimated the oxygen fugacity exponent m to be of 1/6 (Cherniak and Dimanov, 2010;Stimpfl et al., 2005). Klügel (2001) studied the natural zonation of iron-rich orthopyroxene crystals and confirmed the Fe-Mg interdiffusion study extrapolated to higher temperatures by Ganguly and Tazzoli (1994) (Cherniak and Dimanov, 2010). He found DFe-Mg to be equal to 3.10 -19 m 2 /s for T = 1130 °C and QFM buffer redox conditions (Klügel, 2001).

Mg self-diffusion coefficients were determined experimentally in natural orthoenstatites (En88Fs12) by [START_REF] Schwandt | Magnesium self-diffusion in orthoenstatite[END_REF]. Results for the diffusion coefficient are consistent with theoretical values determined by Ganguly and Tazzoli (1994) but the differences may reflect the compositional effect of the Fe 2+ to Mg ratio of the orthopyroxene on the Mg self-diffusion coefficient [START_REF] Schwandt | Magnesium self-diffusion in orthoenstatite[END_REF].

More recently, diffusion experiments between olivine lamellae (Fo30Fa70) and iron-rich orthopyroxene crystals (En90Fs10) have been performed [START_REF] Ter Heege | Experimental Determination of Fe-Mg Interdiffusion Coefficients in Orthopyroxene Using Pulsed Laser Ablation and Nanoscale Thin Films[END_REF]. The results obtained are in fairly good agreement with those of Ganguly andTazzoli (1994). Ter Heege et al. (2006) were also interested in the investigation of the evolution of the diffusion coefficient as a function of oxygen fugacity and found a decrease in the diffusion coefficient with a decrease in oxygen fugacity; the authors estimate an exponent m = 0.15, in agreement with the values found by Ganguly and Tazzoli (1994) and Stimpfl et al, (2005) (1/5.5 < m < 1/6.5) (Cherniak and Dimanov, 2010;Ganguly and Tazzoli, 1994;Stimpfl et al., 2005;[START_REF] Ter Heege | Experimental Determination of Fe-Mg Interdiffusion Coefficients in Orthopyroxene Using Pulsed Laser Ablation and Nanoscale Thin Films[END_REF]. However, the data from Ter Heege et al. ( 2006) are limited and the discussion cannot be taken further.

Recent studies have clarified this expression for the diffusion coefficient as a function of oxygen fugacity (Allan et al., 2017(Allan et al., , 2013) by using Ganguly and Tazzoli (1994) ) in contact with olivines of composition Fo30Fa70 for temperatures between 870-1100 °C at atmospheric pressure and controlling the oxygen fugacity (between 10 -11 -10 -17 Pa) (Dohmen et al., 2016). For the most Enstatite-rich composition, the following Arrhenius equation is found:

𝐷 𝐹𝑒-𝑀𝑔 𝑐 = 1,12.10 -6 𝑓𝑂 2 0.053±0.027 𝑒 -308±23 (III. 23)
Thus, a dependence on oxygen fugacity is identified, which they express according to the relationship:

Log 𝐷 𝐹𝑒-𝑀𝑔 𝑐 = log 𝐷 0 - Q ln(10) RT + n log 𝑓𝑂 2 (III. 24)
Where the activation energy Q = 308 ± 23 kJ/mol, log (𝐷 0 [ 𝑚 2 𝑠 ]) = -5,95 ± 0.83 and n = 0.053 ± 0.027 ~ 1 20 (Dohmen et al., 2016). Comparing this exponent with that found by Allan et al. (2013), the relationship of Dohmen et al. (2016) assumes a minor dependence of the diffusion coefficient on oxygen fugacity (1/20 instead of 1/6 as assumed by Ganguly and Tazzoli (1994)), although that study examined a pyroxene that was very close to enstatite [START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF].

The oxygen fugacity correction applied by Allan et al. (2013) thus appears to overestimate the dependence of the coefficient on fugacity, yielding diffusion coefficients an order of magnitude above the values measured under normal geological conditions (near the NNO; Ni-NiO buffer) (Allan et al., 2013;Dohmen et al., 2016;Ganguly and Tazzoli, 1994). Without these corrections in oxygen fugacity, both parameterizations appear consistent. Ganguly and Tazzoli (1994) explicitly examined a range of orthopyroxenes compositions that includes those in our studies. They studied IW conditions to stabilize Fe-containing crystals and hypothesized a similar fO2 dependence to olivine with an exponent of 1/6, which was incorporated into the work of Allan et al, (2013). The Allan et al. (2013) timescales were updated in a later paper to include the lack of oxygen correction (and also to correct a numerical problem) (Allan et al., 2017).

The subsequent study by Dohmen et al (2016) showed that the dependence of oxygen fugacity on an exponent between 0 and 1/20 would be appropriate from their data, which are limited, not equal to 1/6. Therefore, subsequent work using Ganguly and Tazzoli (1994) took this observation into account and did not apply a correction for oxygen since the experimental data showed that the olivine-like behavior speculated by Ganguly and Tazzoli (1994) did not exist. Similarly, we cannot apply a 1/20 exponent on a dataset where it has never been measured (different compositions of our orthopyroxenes, compared to the En-rich orthopyroxenes of Dohmen et al. (2016)). It seems more reasonable to take this experimental result into account and not include a dependence on oxygen fugacity in our model.

At this time, therefore, there has been no experimental study of the interdiffusion coefficient of an appropriate range of En content other than that of Ganguly and Tazzoli (1994), and no study has examined both the appropriate composition and appropriate ranges of oxygen fugacity.

The influence of oxygen fugacity is therefore still debated and the choice of a certain parameterization of the interdiffusion coefficient depends on crystal compositions and temperature conditions.

Depending on the interdiffusion coefficient used, the timescales are different. Studies have obtained times that differed by a factor of 30 to 40 depending on whether the Allan et al. (2013) or Dohmen et al. (2016) coefficient was used, using the Ganguly and Tazzoli (1994) interdiffusion coefficient also gives shorter times compared to Dohmen et al. (2016) (Fabbro et al., 2018;Flaherty et al., 2018) b. Know the value of the diffusion coefficient and use an appropriate diffusion model (initial condition and boundary conditions) using the mineral-element pair(s) appropriate to its purpose. Particular care must be taken when experimentally determined diffusion data are selected from the literature because the diffusion coefficient they report is only valid under the experimental conditions of the study (T, fO2, P) and cannot be automatically extrapolated to other physical conditions (Solaro-Müller, 2017). c. Obtain a statistical number of diffusion profiles from measuring instruments (EPMA, SEM, LA-ICP-MS, SIMS) allowing to conclude on the timescales involved. d. Know the petrological process on which the diffusion timescale is studied, in order to have a hypothesis on the processes that could be recorded by the crystals and on the thermodynamic variables that have been put into play (in particular the temperature) (Solaro-Müller, 2017).

Problems regarding this strategy mainly concern the diffusion profiles: no preserved profile or profile associated with growth; where the profile has reached a steady state and does not have plateaus; where the profile is too small to be analyzed; where the diffusion coefficient is not available (not for orthopyroxenes) or with a high uncertainty.

These problems can be circumvented by coupling different chronometers (Mg-Sr in plagioclases, Mg-Fe-Al in clinopyroxenes and Mg-Fe in orthopyroxenes to distinguish growth and diffusion, for example) (Costa et al., 2020;Fabbro et al., 2018;Petrone et al., 2018). Indeed, some elements diffusing slowly, their profiles are not supposed to be modified by diffusion and can be used as indicators of initial conditions. This approach has been used in orthopyroxenes by comparing Mg# and Al2O3 content as the profiles are offset and most Mg# gradients are due to diffusion and not growth (Flaherty et al., 2018;Saunders et al., 2012) and is also used in plagioclases (Na-Ca content) (Costa et al., 2003;Morgan et al., 2006) and in olivines (P content) [START_REF] Milman-Barris | Zoning of phosphorus in igneous olivine[END_REF].

III.3.2.2.b. Modelling method

We used the following method, conditions and model, based on the method presented by Allan et al. (2013) and subsequently used by other studies (Couperthwaite et al., 2020;Solaro-Müller, 2017;[START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF]:

a) As the Fe-Mg interdiffusion is frozen during the eruption, according to the shape of the observed sigmoid of the diffusion profiles, it is possible to determine diffusion timescales corresponding to these compositional changes. The acquisition of one-dimensional profiles with EPMA and high-resolution backscattered electron images (BSE) have been acquired in zoned orthopyroxenes with the c-axis in the north-south direction (spatial resolution of 755 nm and analytical resolution close to 0. -Estimation of the temperatures using different techniques: experimental petrology or mineral thermometry (either with mineral pairs, such as magnetite/ilmenite couples or with melt inclusions compositions or mineral/melt pairs (see following chapters) (Putirka, 2008;Sauerzapf et al., 2008)).

-Diffusion coefficient of Ganguly and Tazzoli (1994), which is dependent on orthopyroxene composition, which has been determined in a good compositional range for the crystals in this thesis and without considering the oxygen fugacity. D is extrapolated for our temperatures (no experimental study yet considering both the composition appropriate for our crystals and a good oxygen fugacity range; the coefficient of Dohmen et al. (2016) was not used because Enstatite compositions of crystals are too high compared to the compositions of our orthopyroxenes (see following chapters)).

Diffusion timescales were modelled using a spreadsheet done by D. J. Morgan (now published as a demo version (AUTODIFF) for olivines in Couperthwaite et al. ( 2020)).

Before using the diffusion model, profiles of greyscale values were determined from BSE images using ImageJ and calibrated with EPMA profiles done in the same crystals. Best fit diffusion profiles to the calibrated BSE/EPMA profiles were then modelled using the model spreadsheet (Allan et al., 2013;Couperthwaite et al., 2020;Solaro-Müller, 2017;[START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF], with a finite-difference method that was used to generate a database of simulated diffusion profiles obeying composition-dependent diffusion under a 1D diffusion geometry (Allan et al., 2013;Couperthwaite et al., 2020;Solaro-Müller, 2017;[START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF].

All 1D diffusion profiles are self-similar in time for any specific compositional dependence, for initial boundary conditions (T, fO2, anisotropy and compositional contrasts), the curves are identical in their proportional curvature (Couperthwaite et al., 2020). Thus, if the boundary conditions (fixed) and the initial condition (homogeneous core) are known for a specific measured diffusion profile, the correct diffusion profile shape can be determined from these simulated diffusion curves as the one with the same degree of compositional contrast and asymmetry characteristics. When the appropriate profile has been identified, a linear stretch is applied along the distance axis to fit the data. The spreadsheet measures the distance between the 20 th and 80 th percentile in the measured profile and then scales the appropriate model curve by the correct factor before overlaying them (centered on the 50 th percentile of both curves as a reference point in the centre of the diffusion range) (Allan et al., 2013;Couperthwaite et al., 2020). The stretch factors are used to adjust the known model time to determine the diffusion time, as if for both sample and model (Allan et al., 2013;Couperthwaite et al., 2020):

𝑥 𝛼 √𝐷𝑡 (III. 25)
And

𝑥 𝑠𝑎𝑚𝑝𝑙𝑒 = 𝑛 𝑥 𝑚𝑜𝑑𝑒𝑙 (III. 26)
Where n is the stretch factor, we can write that 𝑡 𝑠𝑎𝑚𝑝𝑙𝑒 is:

𝑡 𝑠𝑎𝑚𝑝𝑙𝑒 = 𝑛 2 𝐷 𝑚𝑜𝑑𝑒𝑙 𝑡 𝑚𝑜𝑑𝑒𝑙 𝐷 𝑠𝑎𝑚𝑝𝑙𝑒 (III. 27)
Before using the images for modelling Fe-Mg interdiffusion, it was verified that the variations in gray level values were only due to variations in Fe-Mg content in the crystals, with no influence of calcium.

Less than 3% of the variations of the gray level in backscattered electrons can be attributed to the variations of calcium (Wollastonite content) ) [START_REF] Ostorero | Time -window into the transcrustal plumbing system dynamics of Dominica (Lesser Antilles)[END_REF]Solaro-Müller, 2017;[START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF]. The Mg number (𝑀𝑔# = 𝑀𝑔 𝑀𝑔+𝐹𝑒 𝑡𝑜𝑡

) is therefore the parameter that seems to best control the grayscale of photos. Morgan, 2010). Other sources of uncertainties concern the parametrization of the diffusion coefficient (accuracy in the determination of the different parameters such as D0, 𝐸 𝑎 ...), but also the accuracy of the measuring instruments used to model the diffusion timescales: the SEM images and EPMA profiles (uncertainties in the calibration of the measurements) (Costa and Morgan, 2010;Morgan et al., 2004;Solaro-Müller, 2017). The resolution of the SEM images can also represent a source of uncertainties (pixel size and its uncertainty, number of lines in the image...), as well as the analysis accuracy of the electron microprobe (step between the points, 1σ uncertainty on the measurement). Combining all these sources of uncertainties, it was shown that, in general, the effects of image quality and resolution or electron microprobe analysis had an influence of only 1% on the calculated timescales and that the largest sources of uncertainties corresponded to the measurement of the diffusion coefficient and temperature (Costa and Morgan, 2010;Solaro-Müller, 2017;[START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF].

The uncertainties are calculated using a Monte Carlo model to calculate the effect of all parameters on the final diffusion times (uncertainty in the diffusion coefficient, electron microprobe measurements and temperature). The uncertainties on temperature, D0 and 𝐸 𝑎 have a logarithmic effect on the uncertainties associated with the diffusion coefficient, thus, the uncertainties on the final diffusion time are asymmetric, with larger values for long times than short times. Indeed, the positive error is larger than the negative error because to calculate them, it amounts to doing, for the positive error: 10 log(𝑡)+log (𝜎) = 10 log(𝑡) 10 log(𝜎) = t σ (III. 28)

With 𝜎, which is the uncertainty on the timescale. And for the negative error:

10 log(𝑡)-log (𝜎) = 10 log(𝑡) 10 log(𝜎) = 𝑡 𝜎 (III. 29)

III.3.3. Diffusion modelling in magnetites

III.3.3.1. Diffusion coefficient estimation in magnetites

Fe-Ti oxides, and magnetite in particular (ideally of formula Fe 2+ Fe 3+ 2O4), are known to rapidly reequilibrate in response to fluctuations in magmatic conditions such as temperature, surrounding fluid composition, oxygen fugacity, pressure, influencing the partitioning of elements at equilibrium (Allan et al., 2017;[START_REF] Aragon | Cation diffusion in titanomagnetites[END_REF][START_REF] Coombs | Reaction rim growth on olivine in silicic melts: Implications for magma mixing[END_REF][START_REF] Devine | Magma Storage Region Processes Inferred from Geochemistry of Fe-Ti Oxides in Andesitic Magma, Soufriere Hills Volcano, Montserrat[END_REF][START_REF] Freer | An experimental study of magnetite-titanomagnetite interdiffusion[END_REF]Nakamura, 1995a;[START_REF] Sievwright | Diffusion and partition coefficients of minor and trace elements in magnetite as a function of oxygen fugacity at 1150 o C. Contrib[END_REF][START_REF] Venezky | Petrology and Fe-Ti oxide reequilibration of the 1991 Mount Unzen mixed magma[END_REF]. They are thus sensitive witnesses of warming episodes in magma storage regions that occur on timescales of days to months prior to eruption of the relevant magmas [START_REF] Devine | Magma Storage Region Processes Inferred from Geochemistry of Fe-Ti Oxides in Andesitic Magma, Soufriere Hills Volcano, Montserrat[END_REF]. Early studies concerning the chemical kinetics of solid solutions of [(Fe3O4)1-x(Fe2TiO4)x] titanomagnetites were carried out at medium temperature conditions (300-600 °C) to investigate the influence of oxidation on the magnetic properties of these minerals [START_REF] Aragon | Cation diffusion in titanomagnetites[END_REF][START_REF] Dieckmann | Defects and Cation Diffusion in Magnetite (IV): Nonstoichiometry and Point Defect Structure of Magnetite (Fe 3 -δO 4 )[END_REF]. The study of the diffusion of Fe, Co, Cr, Al, and Ni cations in magnetites has been investigated by several studies [START_REF] Aggarwal | Point defects and cation tracer diffusion in (TixFe1-x)3-δO4 . II. Cation tracer diffusion[END_REF][START_REF] Dieckmann | Defects and Cation Diffusion in Magnetite (III.) Tracerdiffusion of Foreign Tracer Cations as a Function of Temperature and Oxygen Potential[END_REF][START_REF] Dieckmann | Defects and Cation Diffusion in Magnetite (VI): Point Defect Relaxation and Correlation in Cation Tracer Diffusion[END_REF]. [START_REF] Aragon | Cation diffusion in titanomagnetites[END_REF] studied the intracrystalline diffusion of Ti and Fe at high temperature (1000-1200 °C) as a function of fO2 and spinel composition. More recently, an experimental study determines diffusion coefficients for several elements (Li, Mg, Al, Sc, Ti, Cr, V, Mn, Co, Ni, Cu, Zn, Ga, Y, Zr, Nb, Mo, In, Lu, Hf, Ta, and U) as a function of oxygen fugacity, at 1150 °C and 1 bar [START_REF] Sievwright | Diffusion and partition coefficients of minor and trace elements in magnetite as a function of oxygen fugacity at 1150 o C. Contrib[END_REF]. With its cubic symmetry, there is no variation of diffusivity with crystallographic orientation and the diffusion rates of the various trace elements in this crystal nearly cover four orders of magnitude [START_REF] Sievwright | Diffusion and partition coefficients of minor and trace elements in magnetite as a function of oxygen fugacity at 1150 o C. Contrib[END_REF]. At a given fO2, diffusions coefficients generally increase in this order, for the usual measured between magmatic environments can then be identified, allowing the deciphering of the crystal history (Figure III.22). In this thesis, we applied the CSA approach to opx.

III.5. High-temperature magma mixing experiments

Magma mixing is the chemical interaction between two or more magmas that produces a composition intermediate between the original end-members [START_REF] Bunsen | Ueber die Processe der vulkanischen Gesteinsbildungen Islands[END_REF][START_REF] Jarvis | Crystal and Volatile Controls on the Mixing and Mingling of Magmas[END_REF][START_REF] Wilcox | The Idea of Magma Mixing: History of a Struggle for Acceptance[END_REF]. This phenomenon plays a key role in influencing petrological processes and producing, alongside crystal fractionation, the compositional diversity of igneous rocks [START_REF] Bergantz | Open-system dynamics and mixing in magma mushes[END_REF][START_REF] Jarvis | Crystal and Volatile Controls on the Mixing and Mingling of Magmas[END_REF][START_REF] Morgavi | The Grizzly Lake complex (Yellowstone Volcano, USA): Mixing between basalt and rhyolite unraveled by microanalysis and X-ray microtomography[END_REF]Morgavi et al., , 2013b;;[START_REF] Vogel | Evaluation of magma mixing and fractional crystallization using whole-rock chemical analyses: Polytopic vector analyses[END_REF]). An experimental study was carried out in the Department of Physics and Geology at University of Perugia (Italy) for a study on Kizimen by performing two high-temperature magma mixing experiments (see following chapters).

III.5.1. Starting materials

The experiments consisted in melting two glasses of different compositions in a crucible. Mixing between synthetic or natural endmembers can be undertaken. Synthetic end-members can be created by

Copyrighted material, unauthorised distribution (Figure 4; Kahl et al., 2011) mixing oxide powders following the aimed compositions (silicon dioxide, titanium dioxide, aluminum oxide, iron (III) oxide, magnesium oxide, calcium oxide, sodium carbonate and potassium carbonate).

To do so, after mixing by hand of the oxide powders, a grinding with an agate mortar was done to obtain a fine powder. Then, the two end-members were melted separately inside a high temperature furnace (Nabertherm® HT 04/17) at ambient pressure at 1500 °C in a Pt80Rh20 crucible to produce homogeneous glasses without bubbles or heterogeneities (two melting cycles). To do this, the sample powder was slowly added in a small crucible and then placed in the furnace for a period of time to allow the volatiles time to outgas. The crucible was then gradually filled to the top and left in the oven for 4 hours. Then, the crucible was removed from the furnace and cooled down to room temperature to produce the glass (

III.5.2. Experimental apparatus

The apparatus used for magma mixing experiments was developed at the Department of Physics and Geology (University of Perugia) and presented in details in Morgavi et al. (2015), the Chaotic Magma Mixing Apparatus (hereafter COMMA), which aims to mix viscous natural and synthetic liquids up to a temperature of 1500 °C. The COMMA allows for the initiation of chaotic mixing processes by controlling the mixing kinematics using the Journal Bearing System (JBS) protocol (Morgavi et al., 2015;[START_REF] Swanson | A comparative computational and experimental study of chaotic mixing of viscous fluids[END_REF] ) and ( 2)

the eccentricity ratio with respect to the outer cylinder (𝜀 = 𝛿 𝑅 𝑜𝑢𝑡 = 0.3), where δ is the distance between the two cylinders. The rotations of the two cylinders during the experiment are in opposite directions, thus triggering chaotic mixing (streamlines and velocity are time dependent) [START_REF] Swanson | A comparative computational and experimental study of chaotic mixing of viscous fluids[END_REF]. The device has three main purposes: (a) to allow independent rotations of the inner and outer cylinders at varying speeds, directions, and mixing protocols, (b) to place the apparatus in the hot-spot region of a well-calibrated high-temperature furnace, and (c) finally, to rapidly cool the molten silicates after the mixing protocol has been applied to freeze in time the evolution of the mixing process, avoiding any crystallization of the sample (Morgavi et al., 2015).

Two alumina rods (Al2O3) are attached to the upper and lower axis of the engine to, respectively, serve as the inner cylinder and hold the crucible, with the lower end of the upper alumina rod which is sheathed with a spindle (made of Pt80Rh20) to avoid contamination of the silicate melts during the experiments. As in Morgavi et al. (2015), Pt-PtRh thermocouples (accuracy ± 0.5 °C) are used to measure the temperature within the samples during the experiments (via the alumina rod). The crucible (outer cylinder) is connected to the motor located above the furnace and was placed off-center. The vertical movement of the furnace allows for the preparation of the experimental setup by positioning the outer and inner cylinders outside the furnace. The oven is then brought to the experimental temperature while the sample is still outside the oven (1200 °C for the first experiment). When the rod thermocouple measures the experimental temperature, the experiment is initiated by starting the experimental mixing protocol. The oven is then quickly moved to the final position required to perform the experiment to minimize the time to reach thermal equilibrium (~10-15 min) in the sample. At the end of the experiment, the furnace containing the experimental crucible was quickly moved vertically while keeping the crucible and spindle in their original positions (Morgavi et al., 2015;[START_REF] Rossi | Exponential decay of concentration variance during magma mixing: Robustness of a volcanic chronometer and implications for the homogenization of chemical heterogeneities in magmatic systems[END_REF] (Figure III.23h-i). This final step allows for rapid quenching of the experimental mixture (cooling rate of 41 °C/min) to form a glassy experimental product [START_REF] Rossi | Exponential decay of concentration variance during magma mixing: Robustness of a volcanic chronometer and implications for the homogenization of chemical heterogeneities in magmatic systems[END_REF].

After the experiment, the experimental sample is cored out of the crucible and impregnation in epoxy resin. The sample is then sliced into 2-4 mm thick slices and analyzed by SEM and EPMA to study the textures generated by the mixing process and the spatiotemporal variability of chemical exchange between liquids (see III.2.4).

CHAPTER IV: TIME-WINDOW INTO THE TRANSCRUSTAL PLUMBING SYSTEM OF DOMINICA (LESSER ANTILLES)

Résumé

Il est impossible de prévoir une éruption future mais l'étude détaillée d'éruptions passées pour un système volcanique donné peut permettre de dégager des temps caractéristiques de dynamiques prééruptives pour mieux anticiper le futur comportement volcanique. Ce quatrième chapitre traite de cinq éruptions ponceuses du volcan Morne Trois Pitons-Micotrin de l'île de la Dominique et de la détermination de leurs dynamiques pré-éruptives, la compréhension du comportement des magmas dans le réservoir étant capitale dans la gestion des futures crises volcaniques. La composition des cristaux dépendant de l'environnement magmatique dans lequel ils grandissent (ensemble des conditions de pressions, température et composition du liquide magmatique, par exemple), l'étude des zonations de ces cristaux permet de retracer les changements d'environnement magmatique qu'ils ont subis avant une éruption. Un système de mush transcrustal a été reconnu sous la Dominique (Petites Antilles) avec différentes zones de stockage des magmas qui ont généré la série d'éruptions étudiée : la dernière grande éruption ponceuse « ignimbritique » de ce volcan (Grand Fond -24 ka cal BP) et quatre éruptions Pliniennes de plus faible volume (18-9 ka cal BP). Les dynamiques pré-éruptives des magmas et les changements d'environnement magmatiques subis par les cristaux sont élucidés en combinant la méthode de l'analyse systématique des compositions des cristaux d'orthopyroxènes (zonés et non zonés) avec les temps d'interdiffusion du Fe et du Mg dans ces cristaux.

Deux environnements magmatiques sont reconnus dans le mush magmatique avant ces cinq éruptions, avec des phénomènes de mélanges de magmas qui ont eu lieu, plus ou moins vigoureusement, avant les éruptions successives. De faibles interactions entre les deux environnements ont commencé dans les 10 à 30 ans avant les éruptions Pliniennes de petit volume, mais le mélange soutenu s'est produit au cours des 2 dernières années. Ceci contraste avec les échelles de temps plus longues d'environ une centaine d'années obtenues pour l'éruption plus importante de Grand Fond avec des magmas stockés plus profondément. Il pourrait donc exister un lien entre le temps de réactivation magmatique et les volumes de magmas mis en jeu. En comparant ces temps avec ceux obtenus pour deux autres éruptions ignimbritiques du volcan Morne Trois Pitons-Micotrin et de son voisin le volcan Morne Diablotins, ces temps sont aussi de l'ordre de grandeur de plusieurs décennies (Appendix A). Ces échelles de temps de mélanges de magmas ont des implications significatives pour la gestion des crises volcaniques, avec un signal de réactivation croissant qui pourrait être enregistré à la surface quelques années avant les éruptions. Ces temps pourraient potentiellement permettre de mieux anticiper les délais probables restants avant une éruption. Volcanic eruptions are among the most energetic events on Earth. In a subduction context, eruptions are generally explosive and involve differentiated magmas. These magmas can display variable eruptive styles depending on the volume of magma involved and the relative behaviour of the melt and volatile components during ascent. Deciphering the dynamics of magmas in the crust contributes to a better understanding of magma plumbing system architecture and behaviour, which leads in turn to better management of volcanic crises prior to an eruption 1,2 . Over the last 10 years, physical models suggest that magma chambers are not easy to form and are unstable, contrary to earlier views of these storage areas, with a stable magma chamber dominated by the presence of liquid 3 . Recent works have shown that magmatic reservoirs are likely to be heterogeneously distributed in space and time through the whole crust, as multiple lenses of liquid-dominated magmas, possibly interconnected, surrounded by a mush that is mainly composed of crystals (> 50%) and plutonic rocks derived from the same magmatic reservoirs 4-6 .

Time-window into

Spatiotemporal magma storage and dynamics can be investigated by coupling timescale modelling 7 and the study of the diversity of crystal compositions (in particular zonations), following the Crystal System Analysis approach (CSA) 8 . A CSA approach aims to decipher the record of different pathways experienced by crystals, considering them as sequential of changes in magmatic environments (ME), each defined by constant set of intensive thermodynamic variables such as pressure (P), temperature (T), composition and volatile species fugacities 7,9-12 . It does this by using connectivity diagrams to create a framework in which to organize and interpret the compositional information from crystal cores to the rims 7,9-11 .

Such investigations have successfully been applied on Etna, using olivine crystals to identify magma ponding zones and their varying connections in space and time prior to eruption 7,9-11 . In parallel, following pioneering works on modelling intracrystalline diffusion 13-17 to estimate timescales for changes in magma storage conditions, many applications for diffusion chronometry methods have been found 15,18-29 . Studies have focused on systems including ridge volcanism 13,30 and arc magmatism 16,24,25,31,32 .

Readjustments in mineral-melt equilibria during crystallization can cause a variety of zonation patterns to be formed in growing crystals. These readjustments arise following changes in one or more of the key intensive variables -P, T, oxygen fugacity (fO 2 ), melt composition or volatile abundance 33 . Following the creation of compositional gradients in crystals, intracrystalline diffusion will act to homogenize and re-equilibrate a zoned crystal. Ionic diffusivity in silicates depends on various parameters such as the chemical element, the specific mineral, the crystallographic direction, and the ambient T, P, and fO 2 21,29 process, so the abrupt temperature drop during the eruption leads to effectively instantaneous, cessation of diffusive re-equilibration processes and a "fossilization" of the information linked to pre-eruptive processes in magmatic reservoir. Diffusion chronometry assumes that the diffusion of the elements studied in the mineral must be fast enough to partially erase the zonation but slow enough so that the crystal does not totally equilibrate under the new conditions, allowing temporal information to be recovered 17 .

Our study focuses on Dominica (Lesser Antilles arc), which is the sole island of the arc exhibiting several active volcanic centers and a high magma production rate (Fig. 1a,b). Volcanic activity involving evolved (andesitic-dacitic) magmas is associated with two volcanoes in the central part of Dominica. On Morne Trois Pitons-Micotrin volcanic center, two large pumiceous eruptions (Volcanic Explosivity Index (VEI) of 5 34,35 ) occurred 33 and 24 kyrs ago 34,36,37 . They each generated a fallout layer from a Plinian phase followed by voluminous overflowing pyroclastic density currents sometimes forming welded ignimbritic deposits found in multiple valleys 34 . In the last 20 kyrs, several pumiceous eruptions of smaller magnitude (VEI of 4 34,35 ) have occurred, also originating from Morne Trois Pitons-Micotrin volcanic center. They each generated a Plinian phase with a similar deposit sequence, with dominant pumice fallout followed by low volume of pyroclastic density currents due to collapse of the eruptive column 34 .

In this study, the aim is to examine the spatiotemporal dynamics of the magmas in the plumbing system that fed the successive pumiceous eruptions (one large and four small ones) of the last 24 kyrs of Morne Trois Pitons-Micotrin. These eruptions were chosen because of their recent ages, a good chronostratigraphic control, well-preserved deposits and homogeneous compositions. This permits an investigation of how the storage system of this volcano changes in time and its dynamics, following a larger explosive eruption associated with deeper magma supply (12-16 km 38 ) and smaller eruptions of magmas stored more shallowly (2-8 km 38,39 ). Here, a detailed investigation of orthopyroxene compositions across the last five pumiceous eruptions allows identification of the magma transfer pathways and the existence of disrupting events related to magma remobilisation at depth, by adapting the CSA approach as used for previous large explosive eruptions of Dominica 12 . Timescales associated with the pre-eruptive dynamics were obtained by modelling Fe-Mg interdiffusion between cores and rims 12,25 . Orthopyroxenes of these eruptions are key witnesses of magmatic processes at depth, and are wellsuited for study because they are the best-preserved crystals, with recorded crystal compositions sensitive to temperature variations as determined via experimental works 38,40 . Results are then compared to those of other large explosive eruptions of Dominica, permitting a broader understanding of the transcrustal magmatic system both at this highly active volcanic island and more generally with respect to subduction zone volcanism, which may provide insights useful for volcanic hazard mitigation in Dominica and more generally worldwide.

Previous works

Field study of Dominica eruption products shows that the magma volumes emitted are in the order of 4-5 km 3 Dense Rock Equivalent (DRE) for larger pumiceous eruptions and roughly one order of magnitude lower (< 1 km 3 DRE) for the smaller eruptions 34 .

Furthermore, the storage system beneath Morne Trois Pitons-Micotrin includes two main storage depths 12,38,39 : at ~ 2-8 km depth , based on H 2 O-CO 2 concentrations measured in melt inclusions 39 for the recent small explosive eruptions reservoirs and ~ 12-16 km (300-400 MPa) based on experimental petrology 38 and melt inclusions data 38 for the voluminous pumiceous eruptions reservoirs (Supplementary notes on previous results are available in Supplementary material).

Whole rock compositions of all the studied magmas are andesitic-dacitic, whereas glass compositions are rhyolitic 34,38,39,41 . The pumices possess similar phenocryst abundances of ~ 30 vol% with a dominant proportion of plagioclase (~ 21 vol% for the large eruptions) followed by orthopyroxenes (~ 5% for the large eruptions) and in smaller proportions, clinopyroxenes, Fe-Ti oxides and amphiboles in some units 38,39 .

Results

Orthopyroxene zonation: type and proportions. We focus our study on five eruptions from Morne Trois Pitons-Micotrin 34 : the last voluminous pumiceous eruption of Grand Fond (24 046 ± 357 yr cal BP) and four of the smaller explosive eruptions that occurred in the last 20 kyrs: Goodwill Entrance (Goodwill): 18 095 ± 205 yr cal BP, "Plinian Post Roseau" 1 (PPR1): 12 619 ± 88 yr cal BP, PPR2: 12 513 ± 129 yr cal BP and PPR3: 9 391 ± 94 yr cal BP (Fig. 1a-c and Supplementary Figure S1).

A total of 3224 orthopyroxene (opx) crystals were studied for the five eruptions. Zoned crystals of each eruption were classified according to the type of zonation observed, single zonation (either normal or reverse zonation) and multiple zonation (i.e. Oscillatory zonation) (Fig. 2a-d, Supplementary Table S1). In Grand Fond, of 814 opx studied, only 15% are zoned, with a prevalence in single-zoned and particularly in reverse-zoned crystals compared to normal-zoned ones (Fig. 2e). In the four small explosive eruptions, of the 2410 opx studied, the proportions of zoned crystals are considerably higher (between 51 and 65% of the population being zoned) (Fig. 2f-i). These proportions are considered to be representative and unbiased as the opx were handpicked and mounted randomly before determination of zoning by Scanning Electron Microscopy (see "Material and methods"). In more detail, zoning patterns are more complex for the small explosive eruptions compared to Grand Fond with a higher proportion of multiple-zoned crystals varying between 35 and 61% of opx crystals examined. In the zoned crystals, the proportions of single-zoned vary from 39 to 65% (Fig. 2f-i) with a proportion of normal-zoned from 18 to 68% and reverse-zoned from 32 to 82% (Fig. 2f-i). The multiple-zoned crystals show a dominance of normal zonations followed by reverse zonations (from core to rim) with the exception of Goodwill that shows a higher proportion of reverse then normal zonations (Fig. 2f-i).

Orthopyroxene zonation: compositions. For Grand Fond, 49 core to rim compositional profiles were measured using an electron microprobe micro-analyzer (EPMA): 29 in unzoned crystals and 20 in zoned crys- 60-62 results from the subduction of the northern and southern Atlantic plates beneath the Caribbean plate at a rate of ~ 2 cm/yr 63 . Black: volcanic islands, grey: coral platforms; the map has been modified from a previous study 39 . Bathymetry is in light grey 60 34 . Colors: purple for Grand Fond, blue for Goodwill, yellow for PPR1, red for PPR2, green for PPR3 and grey for the unstudied eruptions. The large explosive eruptions involve 4-5 km 3 of magma Dense Rock Equivalent (DRE); the voluminous pyroclastic density currents, which are several tens of meters thick in the main valleys, are more or less welded. The small explosive eruptions involve magma volumes of at least one order of magnitude lower than the large explosive ones 34 . tals (single-or multiple-zoned). For the four smaller eruptions, 258 profiles have been performed (72 for Goodwill, 59 for PPR1, 61 for PPR2 and 66 for PPR3). All profiles in zoned opx were acquired perpendicular to the zone boundaries and to the long axes of the crystals, far from the corners to reduce the impact of 3-D diffusion effects or recrystallization along the c-axis (Fig. 2b-d) 25,42,43 .

Opx compositions are spread over an En 46-59 compositional range (Figs. 3a1-c5 and 4a1-d5, Supplementary Figure S2). The unzoned opx of the five eruptions have a composition between En 50-56 (Fig. 3a1-a5). Based on the unzoned compositions and the cores compositions of the zoned opx, two populations of zoned opx have S3). The red line shows the location of the electron microprobe profiles (EPMA) and the core, inner ("i") and outer rim ("o") zones are annoted on the multiple-zoned opx (d). From (e) to (i) (drawn with Microsoft Office suite 2019 Version 1808; https:// www. micro soft. com/ fr-fr/ micro soft-365): proportions of the zonations identified in the opx of Grand Fond in the 355-500 µm fraction (e) and in the opx of the small explosive eruptions of the three fractions studied all together (355-500, 250-355 and 125-250 µm fraction, f to i). 4% of the multiple zoned crystals studied had a third rim (less than 5 crystals per eruption). However, this third rim, though existing in the crystal, does not form a plateau in the compositional profiles (< 10 µm), so they will not be considered in the following results (Supplementary Figure S3). These last rims could correspond to very shorts events prior to the eruptions.

been identified (Fig. 3a1-c5): a first population (population 1) whose cores have the same composition as the majority of the unzoned opx (Fig. 4a1-b5) and a second population (population 2) whose core En contents are lower (less magnesian) than the unzoned opx (Fig. 4c1-d5). For Grand Fond, the boundary value separating the two populations is En 53 , En 54 for Goodwill, whereas it is En 52 for PPR1 to .

Population 1 Single-zoned crystals are mostly reverse-zoned, except for PPR1 (normal-zoned) and PPR3 (equal proportions of reverse-and normal-zoned opx) (Fig. 4a1-a5). All the multiple-zoned opx show an initial normal zoning before returning to the initial En content of the cores (Fig. 4b1-b5 and Supplementary Table S2).

Population 2 All single-zoned crystals show a reverse pathway towards the En content domain of population 1 (Fig. 4c1-c5). For multiple-zoned crystals, different pathways are recorded, depending on the eruption (Fig. 4d1-d5). For Grand Fond, they show a reverse pathway, with their inner and outer rims in the same En domain (Fig. 4d1). For Goodwill and PPR2, they show a reverse pathway with outer rims that are richer in En than the inner ones (Fig. 4d2 andd4). For PPR1, only two multiple-zoned opx were analysed and they either show a normal or reverse pathway in the same En domain (Fig. 4d3). The multiple-zoned of PPR3, mostly record a first normal pathway from the core to the inner rims and then a reverse pathway towards the outer rims (Fig. 4d5 and Supplementary Table S2).

Timescales. Timescales have been modelled on the most representative opx zonations. 6 diffusion profiles

were modelled for Grand Fond, as well as 39 for Goodwill, 25 for PPR1, 28 for PPR2 and 35 for PPR3 (see "Material and methods"; Supplementary Figure S3 and Supplementary Table S3). 70% of Grand Fond and PPR1 profiles show the presence of high-frequency bumps and peaks on the slope or/and on the rim side of the compositional profiles, due to a growth component or unstable conditions (Supplementary Figure S3). Modelling of these crystals is not reported in the results due to uncertainty in the zonation origins 12 . Such profiles could be related to fast crystal growth kinetics prior to eruption in the magmatic reservoir, which produce overlapping and complex signals when combined with diffusive relaxation of the initial chemical gradients 44 . Furthermore, 13% of the profiles performed in Goodwill and PPR1-3 (7 for Goodwill, 3 for PPR2 and 4 for PPR3) did not show defined plateaus at the outermost edges of the crystals (Supplementary Figure S4), therefore, only 96 out of the 110 profiles analysed by EPMA in the small eruptions were used for diffusion modelling, with a temperature of 890 °C ("Material and methods", T = 886-895 °C for these eruptions 39 ). For Grand Fond, 6 profiles out of 20 were used for diffusion modelling, with a magmatic temperature of 850 °C12,38, 44 . For Grand Fond, the timescales are distributed between 6 and 66 years, with three timescales around 53 years (Fig. 5a,b; Supplementary Table S3). For the small explosive eruptions, across the diverse zonations considered, Figure 3. Frequency histograms of the En content of unzoned (a1-a5) and zoned opx of the five studied eruptions (single-zoned (SZ) opx: b1-b5 and multiple-zoned (MZ) ones: c1-c5). These frequency histograms highlight two populations of opx: a first population of zoned opx with cores that have the same composition as the majority of the unzoned opx (population 1 in Fig. 4) and zoned opx with cores with lower En contents (less magnesian) than the unzoned opx (population 2 in Fig. 4). The boundary between these two populations are shown with a line on the Figure: En 53 for Grand Fond, En 54 for Goodwill and En 52 for PPR1 to 3. The color for each eruption is the same as in Fig. 2. The magmatic environments (ME) discussed afterwards are specified here for each eruption. n: number of opx analysed.

the timescale estimates show a large variation in both opx populations, from 1 to 32 years. Despite this, the bulk of the analyses represent the final 10 years pre-eruption (Fig. 5c-j and Supplementary Figure S5). For PPR1, timescales are between 2 to 12 years (Fig. 5e,f); Goodwill and PPR2 exhibit a wide timescale dispersion, compared to PPR1 and 3, with few single-zoned opx and some inner rims of multiple-zoned opx that give timescales longer than 20 years, but with the majority of the timescales representing the 20 years pre-eruption (Fig. 5c,d,g,h). The timescale of both inner and outer rims of multiple-zoned crystals are similar to each other and to those of the singly-zoned population. The same observation is made for PPR3, where the timescales go up to 20 years, but again the bulk of the timescales are younger than 10 years (Fig. 5i,j).

Discussion

Considering that the proportion of zoned opx is lower in the larger Grand Fond eruption (15%) than for the four small explosive eruptions (51-65%) (Fig. 2e-i), we can conclude that these latter opx have been more comprehensively impacted by the disrupting events modifying the magmatic environments (ME-as defined in the introduction by specific storage conditions such as temperature, pressure, oxygen fugacity (fO 2 ), volatile content 7,9-12 , see "Material and methods"). Given the proportions of zoned/unzoned opx, our results suggest that the shallow magma ponding zone feeding the recent small explosive eruptions experiences more variable changes between ME than the deep ponding zone experienced during comparable events prior to large explosive eruptions 12 . This difference could be explained by a difference in size of the reservoirs, with Grand Fond's reservoir being larger compared to the four other eruptions, inferred by a difference in the volumes of magmas emitted 34 , as already discussed for another large explosive eruption of the same volcano, Roseau eruption 12 .

Based on the proportions of zoned opx in the five eruptions (Fig. 2), on the proportions of unzoned and cores of zoned opx and their En content (Figs. 3,4), we may define two ME (named ME1 and ME2). We may estimate the relative size of the two ME that can be representative, for Grand Fond, of the volume of the two batches of magma containing respectively ME1 and ME2 (Supplementary Table S4). For Grand Fond, the two ME are ME1 (En 53-56 ) and ME2 (En 49-53 ) (Figs. 3a1,b1,c1 and 4a1-d1). ME1 contains ~ 62% of the opx crystals against ~ 38% for ME2 (Supplementary Table S4). Provided that the crystallinity of the two ME is similar, this would suggest that ME1 is around twice as large than ME2. As with the two other large explosive eruptions studied previously 12 (Roseau from the same volcano and Layou from Morne Diablotins) 12 , reverse-zoned crystals are dominant (88% of the single-zoned crystals, Fig. 2). ME1 contains most of the unzoned opx (59%), few single-zoned cores and multiple-zoned opx (Supplementary Table S4, Fig. 6a). Some multiple-zoned ones have a core and an outer rim in ME1 whereas others have an inner rim in ME1 (for core and outer rim in ME2). ME2 is represented by fewer unzoned opx (26% of the crystals, Supplementary Table S4), cores of single-zoned and multiple-zoned crystals (Supplementary Table S4, Fig. 6a). Some multiple-zoned crystals with cores in ME1 have inner rims in ME2 whereas others have their cores and outer rims in ME2 (and inner rims in ME1) (Fig. 6a). The low variation of En content between these two ME indicates a low temperature difference between them (~ 20 °C, considering an En/T relationship based on experimental data 38 ). It is then likely that these ME were probably located at relatively similar pressure/depths.

For the four recent and small explosive eruptions, opx globally tell us the same kind of story (Fig. 6b-e), with two ME identified for all these eruptions. The relative sizes of the different ME are more or less similar (except for Goodwill, that has a bigger ME2 than for PPR1-3, Supplementary Table S4, Fig. 6b-e), as well as the proportions of zoned and unzoned crystals. Most of the single-zoned crystals have a core in ME2 and a rim in ME1 and, for most of the multiple-zoned crystals, a core and outer rim in ME1 and an inner one in ME2. A large proportion of zoned crystals show low variations of En in the same ME (ME1), they are probably convecting in the same ME. Changes in En content between cores and rims of opx can be interpreted as an evidence of magmatic Figs. 3 and 4,of Goodwill (c,d), PPR1 (e,f), PPR2 (g,h) and PPR3 (i,j). The type of zonation is specified (reverse-zoned or normal-zoned). The timescales are presented as times before the eruptions and can be found in Supplementary Table S3. The colors are the same as in the previous figures. n is the number of profiles modelled (in SZ and MZ opx). The number "n" can be different from the total number of timescales modelled as timescales in MZ opx are sometimes modelled on both the inner and outer rims (e.g.: two timescales modelled in one profile). Individual times and their absolute uncertainties based on the propagation of a temperature uncertainty of ± 25 °C12,38,44 or 9 °C39 and uncertainties associated to SEM images are presented in Supplementary Figure S4. for Grand Fond, Goodwill, PPR1, 2 and 3, respectively. For each eruption, the size of the boxes is approximately proportional to the proportions of unzoned and cores of zoned opx in each ME, using the volumes of ME calculated in Supplementary Table S4. The grey zones show the proportions of unzoned crystals and the white areas, the proportions of zoned crystals (reset to 100%, Supplementary Table S4). The change of ME recorded from the core to the rim of opx is represented with a black arrow for single-zoned (SZ) crystals and a red arrow for multiple-zoned (MZ) crystals. For PPR3 (e), 13 multiple-zoned crystals record a change from ME1 PPR3 to ME2 PPR3 (13 in red). In the case of the studied magmas, an En increase (reverse-zoned) implies a temperature increase, with, or without, magma mixing with a more basic and hotter magma. Conversely, an En decrease (normal-zoned) may be linked to a pressure drop, decrease in melt H 2 O content or mixing with a colder magma 38 . This Figure has been drawn using Adobe Illustrator (Version 24.0.1; https:// www. adobe. com/ fr/ produ cts/ illus trator. html).

processes taking place in magmatic reservoirs such as convection, sedimentation, edge heating or cooling, volatile transfer 4,45-47 leading to changes in crystallization conditions 48 . Following the CSA approach, zoned opx have experienced different ME during magma residence, creating zonations. The CSA diagrams of these four small explosive eruptions illustrate a repeating system with similar or nearly identical behaviors recording a global increase of the temperature (Fig. 6b-e). For these eruptions, the low variation of En content in the opx between the two ME indicates that the two batches of magma containing the two ME are probably located at the same depth in the transcrustal system of Morne Trois Pitons-Micotrin.

Melt inclusions (MI) and matrix glasses have been extensively analyzed in opx 39,41 (Supplementary Figures S6-8; Supplementary Table S5). For all eruptions, MI were abundant and of significant sizes (20-50 µm 39 ) in the cores of the opx but less frequent in the rims and/or too small in size to be analyzed (< 5 µm). MI hosted in opx cores continuously cover a range of composition from 75 to 79% in SiO 2 and 1-2.5 wt% in CaO (Supplementary Figure S8). These small variations in composition do not show significant differences between the two ME (Supplementary Figure S6). These MI were also volatile-saturated 39 . Some matrix glasses were also analyzed showing, for all the eruptions, a SiO 2 content between 75 and 77 wt% comparable to the less silica-rich MI (Supplementary Figures S7-8). The data on major element glass composition (Supplementary Figures S7-8), supports the idea that no basic magma or magma with a significant different composition was involved in the pre-eruptive dynamics of the eruptions. More likely, the ME of this study partook of mixing with magmas of close compositions but possibly different temperatures, to explain the observed crystal zonations and the compositions range of the MI, suggesting a thermally zoned reservoir, with reservoir walls corresponding to a lower temperature environment than its core.

The question then arises of the origin of the unzoned opx (85% of the opx are unzoned for Grand Fond and 35-49% for the four smaller eruptions, Fig. 2). Three possible interpretations are considered. Firstly, unzoned crystals could represent a zero-age population that grew after the mixing event that formed the rims in the zoned crystals. This might be unlikely assuming similar growth rates as unzoned opx should be of similar age as the zoned opx based on their size. Secondly, unzoned crystals might be old crystals that had enough time to fully or near-fully re-equilibrate, but it would take longer than the probable gap between eruptions. Finally, unzoned crystals were present in the magmatic reservoir but in a part of the reservoir that wasn't affected by the mixing process so they do not record that event.

As the majority of the unzoned opx are part of ME1 for all the eruptions, we suggest that they were present before the mixing event. For the unzoned opx with a core composition of ME2, we consider that they represent opx from this particular ME that got mixed with ME1 without recording any changes in compositions. This is particularly true for Grand Fond, with a larger reservoir inferred in which the mixing process was more limited. To better address this question, examining the Al content can be useful as it diffuses slowly in opx and will reflect growth and recrystallisation processes with only a minor diffusion overprint 25,43,49 . Comparison of the Al 2 O 3 content relative to En content for the unzoned and zoned opx has been undertaken for all eruptions (Supplementary Figure S9). In all cases, unzoned opx are in the same compositional domain in Al 2 O 3 (wt%) as the majority of the cores of zoned opx (between 0.4 and 0.8 wt%). Their En content also overlaps the core compositions of the zoned opx. Thus, they were probably in a part of the same source reservoir but a part that never participated in the mixing process. The unzoned phenocrysts are then likely to be representative of the magma state prior to any perturbation (as a mixing event) investigated for diffusion modelling.

The configuration of the magma plumbing system beneath Dominica inferred by previous studies is that of a transcrustal system with batches of magma stored at different depths (~ 12-16 km for the large explosive eruptions and ~ 2-8 km for the small explosive eruptions 38,39 , Fig. 7a,b). Considering these different storage depths, the magmas represented by the different ME are probably stored in a mush. For Grand Fond, as ME2 is also recorded by a large number of opx, we suggest that two batches of magma ME1 and ME2 interacted (Fig. 7b), as for Roseau 12 . They are probably located at the same depth considering the low En variation content in the opx. The large proportion of unzoned crystals (85%) and their abundance in the most voluminous ME1 indicated that a limited mixing process occurred (Fig. 7g,h). For this eruption, the two magmas may probably be stored in two different close reservoirs showing a small composition and temperature difference; the eruption is linked to the injection of magma from ME2 (the smallest reservoir) into ME1, the more voluminous reservoir. The same scenario was described for the large explosive eruptions of Roseau and also for Layou 12 from Morne Diablotins volcano.

For all the small explosive eruptions, the mixing was more efficient, as a larger proportion of crystals are zoned. The volume of magma mobilized for each eruption is lower than for the large explosive eruptions. All these eruptions show the same scenario with proportions of magma from ME1 and ME2 in the same order of magnitude (slightly higher for ME2 in Goodwill). The small En difference between ME1 and ME2 indicates a small difference of temperature and consequently of depth between the two magmas. Considering that (i) a small volume of magma is mobilized for each eruption (less than 1 km 3 ), (ii) the magmas from the two ME have a small composition and temperature difference and are stored at the same depth (shallower than the magma reservoirs of Grand Fond), (iii) the same scenario occurs for all these eruptions in less than 10 kyrs and (iv) compositions of the MI are close and volatile saturation was reached in magmas prior eruption, it is difficult to propose two distinct reservoirs for these eruptions. The most reliable hypothesis is the presence of a thermally zoned reservoir which is not totally emptied after each eruption (Fig. 7a). Between the eruptions, a classical magma evolution within the reservoir would lead to the formation of the two ME, the high-temperature ME1 in the center and low-temperature ME2 being located at the wall of the reservoir. The two magma ponding zones may be linked by replenishment of the shallower one by deeper magmas, which possibly explains the strong similarities in composition between all magmas, except for the volatile content 39,41 (see Supplementary Material).

To further investigate the dynamics of the plumbing system, timescales modelled by intracrystalline diffusion can be coupled with CSA. For Goodwill, as the compositions of the multiple-zoned inner and outer rims are very close, and it is likely that the crystallization of the inner and then outer rims occurred closely in time and could reflect different aspects of the same mixing event. This assumption is verified by the timescales modelled here: timescales modelled on the same multiple-zoned opx for both the inner and outer rims are approximately the same (from 1 to 10 years, Supplementary Table S3 and also for one opx of Grand Fond). Furthermore, for all the small explosive eruptions (Goodwill and PPR1-3), there are no systematic time differences between the inner and outer rims of other multiple-zoned opx and the rims of single-zoned crystals and a greater proportion of them lie within the range 1-10 years. Exceptionally, some rare single-zoned opx give timescales older than 10 years, up to 32 years (Supplementary Table S3). The different magma environments recorded are represented by magma lenses at different depths in a transcrustal system. This schematization agrees with the vision of the magmatic mush showing magma lenses stored at different depths 6 . On top: model of the whole crust with a deep reservoir feeding the large explosive eruptions and a shallower one feeding the small explosive eruptions 38,39 . (a,b) initial stages for small explosive eruptions (Goodwill, PPR1-3) and voluminous pumiceous eruption of Grand Fond; c: rank ordering of the interdiffusion timescales modelled in the opx of Goodwill, PPR1-3 (Fig. 5; Supplementary Table S3). Each curve in (c) can be divided into three gradients (underlined by black arrows): a shallow one (from 32 to 10 years), a steep one (from ~ 10 to 2-3 years), together with that of a pre-eruptive region (timescales from 0 to 2-years, vertical arrow). A steep gradient can indicate a rapid processing whereas a shallow gradient, a much slower processing. (d,e) evolution of the magmas prior to the small explosive eruptions with vigorous magma mixing in a temperature-zoned reservoir, with ME2 being the wall and ME1 the center of the reservoir, in the different periods of time underlined in (c); (f) rank ordering of the interdiffusion timescales modelled in Grand Fond and Roseau 12 , large explosive eruptions from Morne Trois-Pitons Micotrin volcano compared to the small explosive eruptions (Fig. 5; Supplementary Table S3). In (f), for Grand Fond, one gradient can be identified (from 66 to 5 years). For Roseau, two gradients can be identified: a shallow one (from 60 to 7 years) and a steeper one from 7 to 1 year before the eruption. (g,h) evolution of the magmas prior to the large explosive eruption with magma mixing following the period of times shown in (f). T: temperature; black arrow: magma mixing. SZ: single-zoned opx and MZ: multiple-zoned opx. This This could imply that the mixing process evolves through time from single and multiple exchanges, for the small eruptions, 10-32 years before the eruption, to something far more dynamic near the time of eruption (~ 2-10 years) and before the eruption (timescales of ~ 1-year, Fig. 5c-h). For Grand Fond, longer timescales have been obtained (based on the eight timescales modelled on profiles without bumps, Fig. 5a,b; Supplementary Table S3), the remobilization of the magma ponding zone then possibly occurred several decades prior to eruption (up to 66 years before). These timescales estimates highlight a gradual reactivation without an acceleration of the exchanges prior to the eruption.

For Roseau, another large explosive eruption (33 kyrs) in Dominica from the same volcano, presenting the same parameters in terms of storage depth, magma temperature, pressure, fO 2 , shapes and values of the opx zonations, a sustained magma mixing phase has been inferred, 3 to 10 years prior to eruption, after a reactivation period of several decades 12 . As Roseau, some timescales in Grand Fond are also longer than 30 years. The timescales characterizing the processes then seem to be close for both the small explosive eruptions and the larger ones, with longer timescales recorded for Grand Fond.

To decipher the spatiotemporal dynamics in the reservoir, rank ordering diagrams of the timescales modelled for each eruption have been made; curved ramps are associated to each eruption (Fig. 7c,f). These curves can be divided into three gradients for the small explosive eruptions, going from the older timescales to the younger ones with a shallow gradient, then a steeper one and a short vertical one (young population of timescales around 1-2 years) (Fig. 7c). A steep gradient shows a rapid processing whereas a shallow one, a much slower processing. The steep regions of PPR1-3 are much steeper than Goodwill (Fig. 7c); their processing is therefore faster. PPR1's curve only seems to have a steep gradient but this could be biased due to the lower number of crystals studied. This suggests a rapid build up to eruption over a decade at a sustained, relatively high, rate. PPR2 interactions between magmas started about 25 years ahead of the eruption, transitioning to a rapid process some 2-7 years prior to the eruption (Fig. 7c). PPR3 started with a similar slow build up about 20 years pre-eruption, then moved to the rapid process 5 years pre-eruption (Fig. 7c). This all contrasts a bit with Goodwill, which followed a trajectory close to PPR2 but with gradual slopes, and then seems to have stopped about 3 years pre-eruption before suddenly moving to eruption (Fig. 7c).

For Grand Fond, contrary to the other eruptions, Grand Fond's curve seems to be defined by only one gradient: from ~ 70 to 5 years (Fig. 7f). Timescales modelled on bumps go from 1 to ~ 170 years (Supplementary Figure S10, Supplementary Table S3). We conclude that a continuous reactivation is found for Grand Fond, with no rapid build-up in the 10 years pre-eruption as found for the small explosive eruptions.

In comparison, Roseau curve follows the same trajectory as Goodwill, with a slower build up (from 60 years), to a fast one 7 years prior to the eruption 12 (Fig. 7f).

During the decades preceding the small eruptions, some interactions between the two magmas occur and progressively increase up to the last 2 years before the eruptions, where the mixing is more efficient, marked by an increase in the mobilisation rate 2-10 years before all eruptions apart from PPR1 (Fig. 7c), and would gradually leads to overpressure in the reservoir and to the eruption (Fig. 7d,e). Thus, a 10-year warning might be observable e.g., increased seismicity, inflation, degassing and a geochemical response of the hydrothermal system near the Morne Trois Pitons-Micotrin volcano (Valley of Desolation-Boiling Lake system 50 ). These signals could be detectable using seismic data or remote sensing techniques, which is the subject of ongoing research. It is necessary that, during the time separating the eruptions, magmas with a similar composition and probably stored in different pockets in the same order of depth are reinjected in the main reservoir.

Parallel study of plagioclase crystals 38,39 erupted by these five eruptions leads to results consistent with the opx study, with significant development of sieve textures typical of mixing with hotter magma and crystal dissolution (Supplementary Figure S11).

We can then propose a spatiotemporal evolution of the magma plumbing system within a transcrustal system, beneath Morne Trois Pitons-Micotrin volcano in Dominica, using opx as they are trackers of pre-eruptive processes (Fig. 7). Two main magma storage regions are present at Dominica, one at 12-16 km 38 and another at 2-8 km 39 , with variable interactions between batches of magmas through time. 33 kyrs ago, the deeper system saw partial mixing between two magma batches one decade or more prior to the voluminous pumiceous eruptions of Roseau. The eruption of Grand Fond then followed, 9 kyrs later, with similar behaviours but without a dynamic reactivation shortly before the eruption. Since 18 kyrs, magma storage areas, lower in volume, formed at a much shallower depth (~ 2-8 km 39 ), with lower volumes of magma emitted during the successive eruptions of Goodwill Entrance, and PPR1 to PPR3. These eruptions are characterized by vigorous and relatively extensive mixing in a thermally-zoned reservoir, possibly following an injection of a hot magma from a deeper source. These magmatic zones are likely to have been disturbed ~ 10 to ~ 30 years prior to eruption and then more dynamically ~ 2-10 years prior eruption, with activity ramping up prior to the resulting eruption (Fig. 7). Such investigations have a potentially major role in terms of volcanic risk mitigation and management of future explosive volcanic crises.

Material and methods

Sampling. The studied samples are pumices which were collected during field campaigns in Dominica between 2011 and 2013 and were at the base of previous studies 12,34,38,39,41,44 .

Sample preparation.

The unaltered pumice samples were crushed (up to 3 mm). One half was then crushed again into fine powder for whole rock analysis while the other half was sieved into different size fractions (1 mm-750 µm, 750-500 µm, 500-355 µm, 355-250 µm, 250-125 µm). These fractions have been washed in ultrasonic bath and dried at 80 °C for 48 h. They were then observed under the binocular microscope to select the fractions in which the crystals were the more abundant and automorphic. Opx were then handpicked under the binocular microscope from size fractions ranging from 125 to 500 µm. They were mounted in epoxy resin with c-axis in a north-south direction as they have been used for intracrystalline diffusion modelling and polished up to 0.3 µm to the middle part of the opx.

3224 automorphic orthopyroxene (opx) crystals randomly chosen were studied for the five eruptions. Major and trace element studies were performed on opx (Supplementary Figures S2, S6-8). In Grand Fond, 814 opx were studied in the 355-500 µm fraction. In the four small explosive eruptions, 2410 opx have been observed, including 1100 opx in the 355-500 µm fraction, 840 in the 250-355 µm fraction and 470 in the 125-250 µm fraction. The proportions of zoned opx are very similar whatever the considered fraction . Thus, here, the three fractions are gathered and discussed simultaneously.

Plagioclases (plag) and magnetites (mgt) were also handpicked and prepared for textural observations and analyses. Before Scanning Electron Microscope (SEM) or electron microprobe micro-analyzers (EPMA) investigations, selected mounts were all carbon-coated.

Textural observations: scanning electron microscope. All the crystals were observed by using a Scanning Electron Microscope (SEM), Zeiss Supra 55VP (Sorbonne Université, Paris). To identify zoned and unzoned crystals, back-scattered electron (BSE) images were systematically acquired. Proportions of zoned/ unzoned crystals have been determined on all the crystals mounted by looking closely at the SEM images. On zoned crystals, high resolution BSE images were acquired on selected chemical zonations with 20 kV acceleration voltage, a beam current of 8 nA and high integration line (8 integrations per line to reduce the noise) with a dwell time per pixel of 92 µs. These high-resolution BSE images were used for the diffusion modelling and intercalibration with chemical profiles.

Compositional analysis of crystals and melt inclusions by EPMA. Opx crystals have been analysed

for major elements (Si, Na, Ti, Al, Ca, K, Fe, Mg, Mn, P, Cl) with an acceleration voltage of 15 kV, a beam current of 10 nA and a focused beam (CAMECA SX-Five and SX-100; Service Camparis, Paris). Counting time on peak and background for Fe and Mg was set at 80 s and at 10 s for the other elements. The core to rim compositional profiles in zoned opx had a 2 µm step and an average length of 100 µm (~ 4 h per profile). Four points were measured across the unzoned crystals. They were acquired perpendicular to the long axes of the opx and away from the corners to avoid three-dimensional effects such as growth 25,42,43 (Fig. 2b-d).

Glassy melt inclusions in opx and matrix glass were also analysed by EPMA for major elements with the same conditions as described in other studies on Dominica 39,41 . An acceleration voltage of 15 kV was used, with a beam current of 10 nA. Counting time was set at 10 s for the major elements. Na was first analyzed to minimize loss by volatilization under the electron beam 39 . All precautions required for MI investigations were applied: only glassy crystal-and bubble-free ones with no post-entrapment modifications have been investigated 39,41 .

Deciphering spatiotemporal magma dynamics: coupling the crystal system analysis and Fe-Mg interdiffusion modelling of timescales. Stable storage conditions define a given magmatic environment (ME) and are recorded during the growth of the crystals. If the conditions change, the modification of ME is recorded by forming chemical zonations (Supplementary Figure S15) by crystallization, these ones are then equilibrated by diffusion of chemical elements.

A Crystal System analysis (CSA) approach has been developed before on olivines 7,9-11 , according to previous works of this approach 8 . The CSA is the systematic analysis of the crystal compositions to investigate in time and space the changes of crystallization conditions. It is a schematic view of the crystal evolution in different ME that permits to study the conditions of crystal growth from the core to the rim. Each ME is represented by a box and arrows between them indicate the passage, from the core to the rim, of the recorded environments for singlezoned or multiple-zoned crystals. The main path(s) between the ME can then be identified statistically. Here, we applied the CSA approach to the opx 12 . Opx are interesting mineral phases as their composition is sensitive to temperature and thus may help us to track the pre-eruptive dynamics of the plumbing system 38,51 . Compositional plateaus in zoned opx crystals can be identified, with EPMA, as zones of at least 10 µm length where the enstatite content (En content; Mg Mg + Fe ) is constant with a variability of ± 0.5%. If the same compositional plateau is found in a significant amount of crystals, these plateaus have been produced by significative changes in the magmatic conditions 7,52 and can be considered as ME (Supplementary Figure S15). An En increase (reverse-zoned opx) implies a temperature increase, with, or without, magma mixing with a hotter and more basic magma. Conversely, a decrease in the En content (normal-zoned opx) may be linked to a pressure drop, decrease in melt H 2 O content or mixing with a colder magma 38 .

Here, we have coupled the CSA to the interdiffusion Fe-Mg modelling to estimate the timescales of the preeruptive history.

The rate of diffusion of chemical species in the crystal lattice has a strong role on the preservation or resetting of compositional zoning during magma storage at high temperature 52-54 . The diffusion coefficient depends on parameters such as chemical composition (Xi; molar fraction of the mineral constituent element), temperature (T in Kelvin), pressure (P in Pa), oxygen fugacity (fO 2 ) and water fugacity (fH 2 O) 21,29 . Thus, by knowing the diffusion coefficient and by constraining the chemical composition, it is possible to determine Fe-Mg interdiffusion timescales associated to a change in the magmatic conditions that created the zonations in the crystals.

Modelling of the timescales associated to the zonations in the zoned crystals has been done by using the method developed in other studies 25,55 , using the parametrization of the interdiffusion coefficient of Fe and Mg elements in opx 56 , D, which depends on the composition of the opx 12,21 . The Fe-Mg interdiffusion profiles were modelled in one dimension, across the c-axis, parallel to the b-axis of the crystals 25,55 . The oxygen fugacity is an important parameter, though its influence on the Fe and Mg diffusion coefficient is still under discussion.

Previous studies have shown that when integrated in the diffusion coefficient parametrization, differences in the timescales modelled appear. For example, timescales modelling on opx of two Santorini eruptive units give timescales 30-40 times greater than when using another coefficient 25,43,57 .

If the same mechanism of vacancy generation through iron oxidation is taking place, it has been conjectured that the fO 2 dependence could follow the same form as that for olivine (with an exponent of 1/6) 12,56 . This equation parameterization seems to be in accordance with experimental results on opx with Mg# superior to 0.90 57 , arguing for a minor dependence of the diffusion coefficient on fO 2 (with an exponent of 1/20 instead of 1/6) 12 . However, this study was conducted on pyroxene very close to En end member 57 , which is not in adequation for the samples studied here (En 46-59 ).

The correction of the diffusion coefficient for fO 2 made in the previous study 25 seems to overestimate the fO 2 dependence compared to the other diffusion coefficient parametrization 57 , making calculated diffusivities about one order of magnitude higher than the measured values under common geological conditions (close to NNO buffer) 25,44,57 .

Another study in iron rich orthopyroxenes (for temperatures between 1050 and 1200 °C and at Quartz-Fayalite-Magnetite oxygen buffer) reported a diffusivity of D Fe-Mg = 3.10 -19 m 2 .s -1 at T = 1130 °C12, 58 . If the results of the first authors 56 are extrapolated to higher temperatures, they are in good agreement with those of made in the iron rich orthopyroxenes study 12,58 . For these reasons, the oxygen fugacity will not be considered. The expression of the diffusion coefficient we used is valid for a temperature between 500 and 800 °C, an oxygen fugacity of WI at WI + 0.8 log units above the Iron-Wustite buffer (WI: Iron-Wustite) and X Fe = 0.10-0.50, molar fraction of the ferrosilite component 56 . The En range of the opx is En 46-59 (X Fe = 0.45-0.52), which is in the stated calibration range of the parametrization 56 (Supplementary Figure S2). For Grand Fond, the estimated pre-eruptive temperature is 850 ± 25 °C38 . The estimated pre-eruptive temperature of the opx of the four small explosive eruptions is about 890 °C (with an uncertainty of ± 9 °C) 39,59 , which is slightly above the range of used temperatures in the method 21 . Thus, for our diffusion model, the values of the coefficient D used for each profile extrapolate the diffusion coefficient parametrization 56 and are deduced for a temperature of 850 or 890 °C and a given chemical composition (maximum and minimum D calculated by the model using the maximum and minimum En contents of the profile considered).

The initial conditions used for this model are a step-wise profile 12 . Over time, Fe-Mg interdiffusion tends to homogenize the chemical gradient, induced by changes in magmatic conditions, the zonation is then smoothed and the chemical profile is turned into a sigmoid. Since the Fe-Mg interdiffusion is frozen during the eruption, depending on the shape of the sigmoid observed on the profile, it is possible to determine diffusion timescales corresponding to these compositional changes.

The model consists in intercalibrating chemical composition profiles of zoned opx (expressed by the Mg number, Mg# = Mg Mg+Fe tot ) obtained with an electron microprobe with grayscale profiles on high-resolution north-south oriented images taken with SEM (spatial resolution of 755 nm), as the Mg number is the parameter that seems to best control the grayscale of the images 44,55 . These grayscale profiles are produced with the ImageJ software (https:// imagej. nih. gov/ ij/; version 1.52a) and are averaged over a certain width to reduce the noise (between 20 to 50 µm). Following this intercalibration, the profiles are modelled to obtain a time constraint between the possible perturbation recorded by the crystals and the eruption. A database of contrasts of particular compositions is available in our model, the curve best corresponding to the one obtained by intercalibration will be searched in this database 55 . A rescaling between the two curves will have to be performed. This scaling factor depends on several parameters: the ratio between the timescales of the database and the time associated with zoning in the crystal which allows to find the timescale corresponding to the zonation 12 .

The uncertainties associated to the timescales modelled come from the determination of the temperature, the diffusion coefficient and its measurement 19,21 . The accuracy of the measuring instruments used to model the diffusion timescales are also sources of uncertainties: SEM images and electron microprobe profiles (uncertainties on the calibration of measurements, point spacing), as well as the resolution of SEM images. However, the largest sources of uncertainty would come from the measurement of the diffusion coefficient and temperature 12,21,44 (Supplementary Figure S5).

are the same as in the main article (Figure 234). In orange: En contents domain for which the Fe-Mg interdiffusion coefficient parametrization is formulated 1 and other experiments conducted close to the Mg endmember 2 . Note that En46 is just outside the stated calibration range of the interdiffusion coefficient 1 . zonation had not been rebalanced by the Fe-Mg interdiffusion. The black rectangle in (c) represents a portion of the profile affected by the presence of a bump.

(Supplementary Figure S8). In the Al2O3, CaO and FeO versus SiO2 diagrams, the compositions of the MI show clear differentiation trends where compositions become more evolved from plag then cpx and mgt to opx hosted inclusions 4 . The residual glasses also plot in the domain of the melt inclusion compositions 4 . Some additional electron microprobe profile (EPMA) analysis were performed in this study on the MI in the opx of the 355 µm and 250 µm fractions of the small explosive eruptions (Supplementary Figure S6-8, Supplementary Table S4).

Geothermometry and barometry

Grand Fond: The crystallization temperature of Grand Fond's magma has been estimated between 855 and 860°C using titanomagnetites and ilmenites couples; with a fO2 of NNO + 0.6 to + 0.7 3 . Using the opx-melt geothermometer, a temperature range has been calculated, between 870°C to 890°C 11 . The storage pressure of Grand Fond's magmas, using the CO2-H2O ratio content on MI hosted by opx, has been estimated: between 1.6 kbar and 2.7 kbar (corresponding to a depth range of ~7-11 km or 6-10 km considering a crustal density of 2.45 or 2.8 g/cm 3 ), with one inclusion that records a higher pressure of 7.5 kbar (~27 km depth) 10 . The H2O content in these inclusions is between 4.5 to 6.3 wt% 10 . Phase equilibria experiments have been carried out on samples from Roseau ignimbrite coming from the same volcano 3 . They give a pressure between 300 and 400 MPa (~12-16 km considering a mean rock density of 2.45 g/cm 3 or between 11-15 km for 2.8 g/cm 3 ), a temperature of 850±25°C °C, a fO2 of NNO + 1 and ~7 wt% H20.

Small explosive eruptions: the equilibrium temperatures between MI and their host opx have been calculated using the opx-liquid thermometer 11 and are of 895 ± 9°C for Goodwill, 893 ± 9°C for PPR1, 893 ± 10°C for PPR2 and 886 ± 6°C for PPR3 4 . The MI pressure of entrapments have also been estimated between 50 and 200 MPa, which correspond to a depth range of 2-8 km or 1.8-7.3 km (considering a mean rock density of 2.45 g/cm 3 or 2.8 g/cm 3 , respectively), with a melt H2O content up to 3.25 wt% 4 . All arguments tend to consider that magmas from the small explosive eruptions were stored at a maximum pressure of 200 MPa, though part of the crystal assemblage necessarily grew at pressures higher than 200 MPa before ascent and storage in shallower reservoirs. These storage conditions give us the parameters necessary for diffusion modelling.

Supplementary Note: major elements composition of melt inclusions from this study

Some additional EPMA analysis were performed in this study on the MI in the opx of the 355-500 µm and 250-355 µm fractions of the small explosive eruptions (Supplementary Figure S6-8). They have a SiO2 content between 72 and 79 wt% and Na2O+ K2O content that varies from 3.5 to 6.5 wt%. In the K2O, Al2O3, CaO and FeO vs SiO2 diagrams (Supplementary Figure S8), the MI analysed in the opx display a K2O content between 2 to 3 wt%, 11-14 wt% of Al2O3, 1-2.5 wt% in CaO and a FeO content between 2-3.5 wt%, apart from two MI of PPR3 from the 355µm same domain as the majority of the cores of zoned opx (between 0,4-0,8 wt% in Al2O3). Their En content also overlap the core compositions of the zoned opx. f. the core and rims have En contents between En53-56, the ME corresponding to these variations is ME1PPR3 = En52-57. This zoned opx is then a convecting opx in the same ME. By studying the crystals from the core towards the rim, the history of the crystals and the melt containing them in the reservoir before the eruption can thus be reconstructed. Assuming that the growth of the crystal was from the core to the rims, the first ME known by the crystal would therefore be recorded by the core, as it is shown above the EPMA profile, with the evolution of a crystal in time. This

Eruption

Grand

CHAPTER V: CRYSTALS RECORD OF MAGMA DYNAMICS BEFORE THE 1956 FLANK COLLAPSE ERUPTION OF BEZYMIANNY Résumé

Dans ce chapitre, l'éruption ciblée est la phase paroxysmale de la première éruption historique du volcan Bezymianny (30 mars 1956) au Kamchatka [START_REF] Braitseva | Bezymianny: eruptive history and dynamics[END_REF][START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF][START_REF] Mania | Deformations and Morphology Changes Associated with the 2016-2017 Eruption Sequence at Bezymianny Volcano, Kamchatka[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]. Après une période de repos de 1000 ans, une déstabilisation du flanc du volcan à grande échelle (0,5 km 3 , 30 km de distance parcourus par l'avalanche de débris générée) a entraîné la dépressurisation rapide d'un cryptodôme stocké dans le flanc du volcan et une forte explosion dirigée latéralement le 30 mars 1956 [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF]. Ce blast a détruit la partie orientale de l'édifice volcanique, ainsi qu'une zone de 500 km 2 autour du volcan, produisant ~0,2 km 3 de dépôts. Il a été suivi quelques minutes après par une phase Plinienne (colonne éruptive de 34-38 km de haut) générant des retombées de cendres et de ponces ainsi que des courants de densité pyroclastiques ponceux par écroulement de la colonne [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF]. L'intérêt de l'étude de cette éruption est donc de déterminer les temps correspondants à des processus magmatiques sur ce volcan qui n'a pas connu d'éruptions depuis 1000 ans ainsi que les temps de réactivation, de remontée du magma et de son séjour dans le cryptodôme.

Ces explosions latéralement dirigées sont des événements explosifs rares mais très dévastateurs consécutifs à un effondrement sectoriel majeur d'un volcan, avec le potentiel de dévaster une zone de plusieurs centaines de km 2 autour du volcan, en raison de puissants courants de densité pyroclastiques dilués et turbulents, comme ce fut le cas lors de l'éruption du 18 Mai 1980 du Mont Saint Helens aux Etats-Unis, dont le blast a dévasté 600 km 2 sur le flanc nord du volcan [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF][START_REF] Lipman | The 1980 eruptions of Mount St[END_REF].

En combinant les chronomètres pétrologiques des orthopyroxènes et des magnétites provenant d'échantillons des différentes phases de cette éruption, nous avons montré que des mouvements magmatiques se sont produits au cours des 2 à 3 dernières années avant l'éruption, suivis d'une ascension du magma dans un cryptodôme au cours des six derniers mois. Le magma contenu dans le cryptodôme a subi un phénomène de dégazage et/ou un processus de réchauffement deux mois avant l'éruption, formant des zonations inverses dans les orthopyroxènes et dans les magnétites. Les autres cristaux de plagioclases et d'amphiboles contenus dans ce magma, ont également enregistré ce processus de réchauffement. Le magma générant la phase Plinienne montre un processus de réchauffement encore plus court dans le réservoir magmatique. Des modifications de la température, de la fugacité de l'oxygène (équivalent de la pression partielle en oxygène dans le magma) ou de la paramétrisation du coefficient de diffusion ont été testées afin d'étudier leurs effets sur les échelles de temps modélisées dans les cristaux d'orthopyroxènes. Une augmentation de la température entraîne des temps plus courts (l'inverse est observé pour une température plus basse). De plus, des temps plus courts sont observés si la fugacité d'oxygène est prise en compte. Un changement du coefficient d'interdiffusion Fe-Mg double les échelles de temps modélisées (Dohmen et al., 2016).

Il est donc important de comprendre ces dynamiques pré-éruptives avant des éruptions où la période de réactivation sismique peut être relativement courte. Dans le cas de l'éruption de 1955-1956, la présence d'une seule station sismique située à 42 km, laisse place à un questionnement par rapport aux signaux sismiques détectés ; il est probable que la totalité des signes précurseurs n'aient pas vraiment tous été enregistrés [START_REF] Tokarev | The prediction of large explosions of andesitic volcanoes[END_REF].

Crystals record of magma dynamics before the 1956 flank collapse eruption of Bezymianny

Provisional list of authors: Lea Ostorero 1* , Georges Boudon 1 , Hélène Balcone-Boissard 2 , Thiebaut d'Augustin 2 (rest of the author list to be defined) 1 Université Paris Cité, Institut de physique du globe de Paris, CNRS, F-75005 Paris, France

V.1. Introduction

In volcanology and more specifically, in volcano monitoring, the issues at stake are the uncertainties regarding the time of onset of an eruption, from the appearance of signs of volcanic unrest, to the eruption. The time window between these two phenomena are a major field of research, as it can potentially pose a threat to citizens lives and property and have an impact on them on a large scale.

Eruptions of arc volcanoes are all the more dangerous due to the compositions of the magmas and their differentiation with time. The compositions of the magmas are generally intermediate to silicic, volatilerich, and with the potential of experiencing transitions between eruptive styles, giving to the eruptions a more unexpected behavior to monitor. In explosive eruptions, a type of eruption that was only recently understood is the directed volcanic blasts. These eruptions are powerful explosions with a significant laterally-directed component, that can generate devastating, high energy pyroclastic density currents [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF][START_REF] Glicken | Rockslide-debris avalanche of May 18[END_REF][START_REF] Voight | Time scale for the first movements of the May 18 eruption[END_REF][START_REF] Voight | Catastrophic rockslide avalanche of May 18[END_REF]. Despite their relatively small volumes of erupted magma (< 1 km 3 ), these eruptions lead to complete devastation in areas of up to hundreds of square kilometers. The name of these eruptions (directed blast) was born following the 1956 flank collapse eruption of Bezymianny in Kamchatka, after the observations from a long distance by [START_REF] Gorshkov | Gigantic eruption of the volcano bezymianny[END_REF]. This eruption was misinterpreted at the time and it was not until the eruption of Mount St Helens in 1980 [START_REF] Glicken | Rockslide-debris avalanche of May 18[END_REF][START_REF] Voight | Time scale for the first movements of the May 18 eruption[END_REF][START_REF] Voight | Catastrophic rockslide avalanche of May 18[END_REF] and the detailed study of Bezymianny eruption by [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF] that the phenomenology of the eruption was fully understood.

The combined growth of a lava dome at the top of the volcano and the injection of a part of the magma in the form of a cryptodome in the eastern flank of the volcano caused an extreme deformation of its flank, that required approximately 2-6 months [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF], until its destabilization generating a debris avalanche and forming an open horseshoe-shaped structure [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF][START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF][START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF][START_REF] Belousov | DEBRIS AVALANCHE OF THE 1956 BEZYMIANNY ERUPTION[END_REF]. The sudden depressurization of the magma contained in the cryptodome caused an extremely devastating laterally directed blast, destroying the entire eastern flank of the volcano, with trees all knocked down in the same direction over several hundred km 2 [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF][START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF]. As at Mount St Helens in 1980 [START_REF] Criswell | Chronology and pyroclastic stratigraphy of the May 18, 1980, Eruption of Mount St. Helens, Washington[END_REF][START_REF] Lipman | The 1980 eruptions of Mount St[END_REF], a Plinian phase immediately followed the laterally directed blast [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF].

Understanding the magma dynamics and the timescales of the magmatic processes prior to this kind of eruptions is then of particular interest. Information trapped in volcanic rocks emitted by the eruption are of different types: crystals and diffusion timescales to date magmatic processes, melt inclusions to understand pre-eruptive magmatic storage conditions and residual glasses to inform on decompression processes during the eruption. We will here focus on crystals, which are powerful petrological tools as they can be seen as petrological clocks, activated when storage conditions change [START_REF] Anderson | Probable relations between plagioclase zoning and magma dynamics, Fuego Volcano, Guatemala[END_REF][START_REF] Giacomoni | Plagioclase as archive of magma ascent dynamics on "open conduit" volcanoes: The 2001-2006 eruptive period at[END_REF]Petrone and Mangler, 2021;[START_REF] Streck | Mineral Textures and Zoning as Evidence for Open System Processes[END_REF]. These conditions can change following a magma injection or magma mixing processes, for example, as well as crystallization and differentiation processes (Costa, 2021;Costa et al., 2020;[START_REF] Hawkesworth | Time scales of magmatic processes[END_REF]Petrone and Mangler, 2021). After a magmatic perturbation, rims are created in crystals (called "zonations"), with different compositions. Multiple-zoned crystals with complex textures can appear, showing the complexity of the magmatic processes recorded. These zonations can help to decipher the different pathways of crystals that grew in the plumbing system and highlight changes in magmatic environments (ME; defined by changes in temperature, pressure, oxygen fugacity (fO2) and volatile element abundance) (Kahl et al., 2011;[START_REF] Ubide | Deep Magma Storage Revealed by Multi-Method Elemental Mapping of Clinopyroxene Megacrysts at Stromboli Volcano[END_REF]. The timescales of magmatic processes then have a large variability, depending on the magma dynamics into the magmatic reservoir and they are mainly controlled by physical parameters such as magma temperature, rheological properties of the environment and magmatic flows in the magmatic system. Residence timescales of crystals and magma dynamics of the reservoir can play a role on magma viscosity, crystallization rate and exsolution rate of volatile elements, influencing the future eruptive style [START_REF] Gudmundsson | Magma chambers: Formation, local stresses, excess pressures, and compartments[END_REF].

Combining diffusion timescales results from multiple mineralogical chronometers can allow to obtain several independent timescale estimates to have important temporal information on the pre-eruptive magmatic processes (Chamberlain et al., 2014;Fabbro et al., 2018). Few studies have modelled diffusion in multiple minerals from the same eruption [START_REF] Brugman | Common assumptions and methods yield overestimated diffusive timescales, as exemplified in a Yellowstone post-caldera lava[END_REF]Chamberlain et al., 2014;[START_REF] Cooper | Rapid assembly and rejuvenation of a large silicic magmatic system: Insights from mineral diffusive profiles in the Kidnappers and Rocky Hill deposits, New Zealand[END_REF]Fabbro et al., 2018;Flaherty et al., 2018;[START_REF] Magee | The Lead-up to Mount Etna's Most Destructive Historic Eruption[END_REF][START_REF] Shamloo | Decadal transition from quiescence to supereruption: petrologic investigation of the Lava Creek Tuff, Yellowstone Caldera[END_REF]Singer et al., 2016) or modelled the timescale given by more than one compositional boundaries (Druitt et al., 2012;Kahl et al., 2011;Petrone et al., 2018Petrone et al., , 2016;;[START_REF] Rout | Constraints on non-isothermal diffusion modeling: An experimental analysis and error assessment using halogen diffusion in melts[END_REF].

Bezymianny and its 1956 flank collapse eruption is our case study as this catastrophic eruption provides an opportunity to investigate the changes that occurred in the magmatic system in the years preceding the flank collapse through petrological investigations. A preclimactic phase began in October 1955, preceding the climactic and devastating phase of the flank collapse on 30 March 1956 [START_REF] Gorshkov | Gigantic eruption of the volcano bezymianny[END_REF][START_REF] Gorshkov | Bezymianny Volcano and Peculiarities of Its Last Eruptions in[END_REF].

We then aim to answer the two main following questions:

(1) Which magmatic processes occurred before the catastrophic eruption of Bezymianny that led to the laterally directed blast and the Plinian phase? (2) On which timescales? In our approach, we combine the analysis of the pre-eruptive timescales of two different types of minerals, orthopyroxenes and magnetites, using geothermometry to also estimate the temperature conditions of magmatic storage. We propose a space and time revisited scenario of the pre-eruptive magma dynamics.

V.2. Geological context and the 1956 flank collapse of Bezymianny

The "Pacific Ring of Fire" is a long chain of 40 000 km long with more than 450 active volcanoes that produce explosive and effusive eruptions, with Kamchatka being an active region of this chain [START_REF] Eichelberger | Preface, in: Volcanism and Subduction: The Kamchatka Region[END_REF]. Kamchatka Peninsula is located in the Kuriles-Kamchatka arc, in the northwestern Pacific, where the Pacific plate subducts under the Kamchatka-Okhotsk continental block and the associated volcanism results in violent explosive eruptions or effusive eruptions of andesiticdacitic/andesitic magmas, as well as basaltic magmas and also some rhyolitic ones [START_REF] Dobretsov | Migration paths of magma and fluids and lava compositions in Kamchatka[END_REF]

(Figure V.1a).
The rate of oblique subduction is fast under Kamchatka (~8 cm/yr) [START_REF] Demets | Oblique convergence and deformation along the Kuril and Japan Trenches[END_REF], where the oceanic Pacific plate has a Cretaceous age (90-120 Ma) [START_REF] Dobretsov | Migration paths of magma and fluids and lava compositions in Kamchatka[END_REF]. Three volcanic zones of Neogene-Quaternary age are present in Kamchatka: the Sredinny Range (Central-Kamchatka belt, mainly with basaltic volcanic deposits and some andesites and dacites) and an Eastern zone subdivided in two: the Central Kamchatka Depression (CKD) and the Eastern Volcanic Front (southern part of Kamchatka to the northern Kuriles), where all the historical eruptions occurred [START_REF] Ponomareva | Late Pleistocene-Holocene volcanism on the Kamchatka Peninsula, Northwest Pacific Region, in: Volcanism and Subduction: The Kamchatka Region[END_REF]

(Figure V.1a).
The CKD volcanoes are the most magma-producing arc volcanoes on Earth (as they lie above the slab edge) [START_REF] Braitseva | Ages of calderas, large explosive craters and active volcanoes in the Kuril-Kamchatka region, Russia[END_REF]Levin et al., 2002) and are the most studied in Kamchatka [START_REF] Belousov | Pyroclastic surges and flows from the 8-10 May 1997 explosive eruption of Bezymianny volcano, Kamchatka, Russia[END_REF][START_REF] Belousov | DEBRIS AVALANCHE OF THE 1956 BEZYMIANNY ERUPTION[END_REF]Davydova et al., 2018[START_REF] Davydova | Petrology of mafic enclaves in the 2006-2012 eruptive products of Bezymianny Volcano, Kamchatka[END_REF][START_REF] Dorendorf | Hydrated sub-arc mantle: a source for the Kluchevskoy volcano, Kamchatka/Russia[END_REF][START_REF] Journeau | Seismic tremor reveals active trans-crustal magmatic system beneath Kamchatka volcanoes[END_REF][START_REF] Kersting | Klyuchevskoy Volcano, Kamchatka, Russia: The Role of High-Flux Recharged, Tapped, and Fractionated Magma Chamber(s) in the Genesis of High-Al2O3 from High-MgO Basalt[END_REF][START_REF] Khubunaya | The mineralogy of the Klyuchevskoi magnesian basalts depicting the fractionation in the magma chamber[END_REF][START_REF] Koulakov | Rapid changes in magma storage beneath the Klyuchevskoy group of volcanoes inferred from time-dependent seismic tomography[END_REF][START_REF] Koulakov | Mantle and Crustal Sources of Magmatic Activity of Klyuchevskoy and Surrounding Volcanoes in Kamchatka Inferred From Earthquake Tomography[END_REF][START_REF] Koulakov | Three different types of plumbing system beneath the neighboring active volcanoes of Tolbachik, Bezymianny, and Klyuchevskoy in Kamchatka[END_REF][START_REF] Mania | Deformations and Morphology Changes Associated with the 2016-2017 Eruption Sequence at Bezymianny Volcano, Kamchatka[END_REF][START_REF] Mironov | Final Stages of Magma Evolution in Klyuchevskoy Volcano, Kamchatka: Evidence from Melt Inclusions in Minerals of High-Alumina Basalts[END_REF][START_REF] Ozerov | The evolution of high-alumina basalts of the Klyuchevskoy volcano, Kamchatka, Russia, based on microprobe analyses of mineral inclusions[END_REF]Portnyagin et al., 2007b[START_REF] Portnyagin | Transition from arc to oceanic magmatism at the Kamchatka-Aleutian junction[END_REF][START_REF] Shcherbakov | Phase equilibria constraints on preeruptive magma storage conditions for the 1956 eruption of Bezymianny Volcano, Kamchatka, Russia[END_REF]Volynets et al., 1999b) Bezymianny (2886 m, above sea level), in the Klyuchevskoy volcanic group, is a dome-building stratovolcano dating from the Pleistocene period ( 10.5-11 ka BP) [START_REF] Braitseva | Bezymianny: eruptive history and dynamics[END_REF][START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF][START_REF] Mania | Deformations and Morphology Changes Associated with the 2016-2017 Eruption Sequence at Bezymianny Volcano, Kamchatka[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]. Bezymianny volcanism began with the formation of a cone on the southern flank of the presumably extinct Kamen volcano, which, through eruptions, allowed the building of a stratovolcano [START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF] (Supplementary Figure V.1). Between 2.4-1.7 ka and 1.35-1 ka BP, major eruptive activity occurred at Bezymianny [START_REF] Braitseva | Bezymianny: eruptive history and dynamics[END_REF][START_REF] Mania | Deformations and Morphology Changes Associated with the 2016-2017 Eruption Sequence at Bezymianny Volcano, Kamchatka[END_REF]. The pre-historic eruptive activity formed lava flows, extrusive lava domes and extensive pyroclastic density current deposits, with magma compositions ranging from basalts to andesites, with volcanic activity preceded by long phases of dormancy [START_REF] Bogoyavlenskaya | Bezymianny volcano[END_REF][START_REF] Braitseva | Bezymianny: eruptive history and dynamics[END_REF][START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]. Before Bezymianny recent volcanic activity, its summit was at 3085 m and had a 500 m wide ill-defined crater, on the summit, practically filled by flank extrusive viscous lava domes, cut by a deep gorge descending from the western rim of the crater [START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF][START_REF] Gorshkov | Gigantic eruption of the volcano bezymianny[END_REF]. The last repose period of this volcano was from 1000 years BP to 1955 AD [START_REF] Braitseva | Bezymianny: eruptive history and dynamics[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF].

V.2.1. 1956 flank collapse

After 23 days of intense seismic swarm beginning on 29 September 1955 (hypocenters estimated at less than 5 km from the floor of the active crater), the first historical eruption of Bezymianny occurred on 22 October 1955 (Supplementary Figure V.2) [START_REF] Gorshkov | Gigantic eruption of the volcano bezymianny[END_REF][START_REF] Tokarev | The prediction of large explosions of andesitic volcanoes[END_REF][START_REF] Tokarev | Volcanic Earthquakes of Kamchatka[END_REF]. It began by a preclimactic phase with an explosive activity (Vulcanian character) bursting out of a new crater, with frequent ash explosions with plumes sometimes rising more than 6 km high [START_REF] Belousov | DEBRIS AVALANCHE OF THE 1956 BEZYMIANNY ERUPTION[END_REF][START_REF] Gorshkov | Gigantic eruption of the volcano bezymianny[END_REF]. Eruptive activity diminished by 21 November, with rarer and weaker eruptions (plumes < 1.5 km) and after the end of November, a lava dome may have grown slowly in the new crater (seen for the first time on 25 January 1956)

(Supplementary Figure V.2).
Simultaneously, a batch of magma probably intruded the eastern flank of the volcano (cryptodome), as the southeastern slope of the cone was slowly uplifted (high as 100 m, photographs in February 1956) [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF][START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF]

(Supplementary Figure V.
2). This large deformation, that accompanied the cryptodome formation in the upper part of the volcanic cone, contributed to the structural instability of the edifice but Bezymianny was able to accommodate the deformations plastically for a long time, allowing magma volumes to be accumulated (dome/cryptodome volume of 0.15-0.2 km 3 ) before the final rupture [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF].

Up to March 1956, weak explosions with ashes plumes took place (up to 3 km high) (Supplementary Figure V.2). On 30 March 1956, the climactic phase of the eruption occurred unexpectedly during a general weakening of the volcanic eruption and seismic activity. The collapse of the highly deformed southeastern flank of the volcano generated a debris avalanche of 0.5 km 3 that rushed down at a speed of ~60 m/s, and covered an area of 36 km 2 , leading to the opening of a large horseshoe-shaped crater of 1.7x2.8 km, to the southeast [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF][START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF][START_REF] Belousov | DEBRIS AVALANCHE OF THE 1956 BEZYMIANNY ERUPTION[END_REF]. Unroofing the cryptodome, the flank-collapse triggered a catastrophic laterally directed blast towards the southeast [START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF]

(Supplementary Figure V.2).
The diluted and turbulent pyroclastic density current destroyed an area of ~500 km 2 , with a velocity of more than 100 m/s (< 250 °C). The volume of the deposits was estimated at 0.2-0.4 km 3 [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF][START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF]. The kinetic energy of the blast was of 1.2.10 17 J (compared to 10 14 -10 15 J for 1980 Mount St. Helens) [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF][START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF]. Following this explosion, the depressurization of the conduit triggered a Plinian phase that generated a 34-38 km high eruptive column, which left an extensive airfall deposit (0.2-0.3 km 3 ) to the north and ash and pumiceous pyroclastic density currents with a volume of 0.5 km 3 [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF][START_REF] Belousov | DEBRIS AVALANCHE OF THE 1956 BEZYMIANNY ERUPTION[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]

(Figure V.1b; Supplementary Figure V.1-2).
The total volume of the eruptive products was between 0.9-1.2 km 3 (blast deposits, Plinian fallout and pyroclastic density currents from the Plinian column) [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF]. The base of the new crater was approximately at 2250 m after the eruption (height reduction over vent of 750 m) [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF]. The 1955-1956 eruption was immediately followed by an open-conduit dynamic, with the quasi continuous growth of lava domes, periodic dome collapses and associated block and ash pyroclastic density currents until 1976 (Supplementary Figure V.2) [START_REF] Davydova | Petrological evidence of rapid evolution of the magma plumbing system of Bezymianny volcano in Kamchatka before the December 20th, 2017 eruption[END_REF][START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]. Lava dome growth stopped since the 1970s and the eruptive style changed to short-lived strong Vulcanian explosive eruptions once or twice per year and extrusive-explosive or effusive periods [START_REF] Davydova | Petrological evidence of rapid evolution of the magma plumbing system of Bezymianny volcano in Kamchatka before the December 20th, 2017 eruption[END_REF][START_REF] Girina | Chronology of Bezymianny Volcano activity[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]. The volcano remained dormant between September 2012 and December 2016 but this period of dormancy ended with the onset of an effusive eruption with lava dome growth (December 2016) [START_REF] Mania | Deformations and Morphology Changes Associated with the 2016-2017 Eruption Sequence at Bezymianny Volcano, Kamchatka[END_REF]. Since then, seven explosive eruptions have occurred (three in 2017, two in 2019, one in 2020 and one in 2022) [START_REF] Davydova | Petrological evidence of rapid evolution of the magma plumbing system of Bezymianny volcano in Kamchatka before the December 20th, 2017 eruption[END_REF]Global Volcanism Program, 2022).

V.2.2. Magma plumbing system architecture prior to the 1956 flank collapse and climactic phase

Several studies have estimated the pre-eruptive storage conditions of the magmas and have mainly found two reservoirs (but estimated at different depths, depending on the studies). By studying crystallization conditions of phenocrysts, they were estimated at 11-30 km (300-800 MPa, 875-925 °C and > 6 wt% H2O) [START_REF] Kadik | Physicochemical Crystallization Conditions and Genesis of Andesites[END_REF], agreeing with [START_REF] Almeev | Problems of the Stoichiometry and Thermobarometry of Magmatic Amphiboles: An Example of Hornblende from the Andesites of Bezymyannyi Volcano, Eastern Kamchatka[END_REF], reporting a pressure of ~600 MPa (23 km) and T of 800-920 °C (for a density of 2700 kg.m -3 ). These pressures also agree with [START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF], who studied amphibole reaction rims development in cryptodome fragments, as well as phase equilibria simulations (Almeev et al., 2013a(Almeev et al., , 2013b)). From experimental petrology and reaction rims in amphiboles, prior to 1956 eruption, a depth of ~8 km was also evidenced (200 MPa, 775 °C for amphibole, plagioclase, orthopyroxene assemblage, for a density of 2700 kg.m -3 ) [START_REF] Shcherbakov | Phase equilibria constraints on preeruptive magma storage conditions for the 1956 eruption of Bezymianny Volcano, Kamchatka, Russia[END_REF]. These authors also inferred that the magma leading to the Plinian phase was last and transiently stored at least during 40 days between ~2-4 km (50-100 MPa, 890-930 °C). The eruptive activity from 1956 to 2012 showed a reverse acid-to-basic trend (from relatively silicic magma in 1956: andesite to basaltic andesite) [START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF], which could be explained by a multilevel magma plumbing system, with at least two magma reservoirs at different depths. From amphibole thermobarometry, [START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF] indicated that these reservoirs would be between [8][9][10][11] MPa at 975 ± 25 °C. These two last pressure ranges correspond well to the crustal reservoirs characterized by seismic investigations at ~7 km and 18 km [START_REF] Fedotov | The magmatic system of the Klyuchevskaya group of volcanoes inferred from data on its eruptions, earthquakes, deformation, and deep structure[END_REF][START_REF] Koulakov | Three different types of plumbing system beneath the neighboring active volcanoes of Tolbachik, Bezymianny, and Klyuchevskoy in Kamchatka[END_REF][START_REF] Ter Heege | Experimental Determination of Fe-Mg Interdiffusion Coefficients in Orthopyroxene Using Pulsed Laser Ablation and Nanoscale Thin Films[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]. Furthermore, it was shown that Bezymianny, Klyuchevskoy and Tolbachik volcanoes derive their fluids from a common magma chamber at 30 km depth [START_REF] Koulakov | Three different types of plumbing system beneath the neighboring active volcanoes of Tolbachik, Bezymianny, and Klyuchevskoy in Kamchatka[END_REF][START_REF] Shapiro | Deep and shallow long-period volcanic seismicity linked by fluid-pressure transfer[END_REF]. Eruption triggering could result of regular injections of hotter magma saturated in volatiles, in the shallow chamber, maintaining a stable temperature (940-980 °C) [START_REF] Davydova | Petrology of mafic enclaves in the 2006-2012 eruptive products of Bezymianny Volcano, Kamchatka[END_REF].

V.3. Methods

V.3.1. Sampling on the field

Samples were collected in the laterally directed blast (called "blast" further down) and in the pumiceous pyroclastic density currents deposits from the Plinian phase of the 1956 climactic phase, during a threeweek field campaign which took place in Kamchatka in August 2019. In the blast, samples with different lithologies, sorting and grading were collected, resulting from the sedimentation regimes of the diluted and turbulent pyroclastic density current [START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF]. Samples coordinates are given in Table V. 1. Supplementary Figure V.1 shows one of the sampling sites of the blast deposits that consist of three layers (layers A, B and C) [START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF].

Table V.1: Description of the 1956 climactic phase samples of Bezymianny studied here.

Localization, coordinates, nature of the samples, as well as the density and vesicularity measured for the individual clasts studied here are specified, following the method presented in the Methods section. The density and vesicularity ranges of all the fragments are specified (Bezy 3, Bezy 8 and Bezy 9).

V.3.2. Sample preparation

The unaltered andesitic samples were crushed (up to 3 mm). One half was then crushed again into fine powder for whole rock analysis while the other half was sieved into different size fractions (1 mm-710 µm, 710-500 µm, 500-315 µm, 315-250 µm, 250-125 µm) for the crystal studies. These fractions have been washed in ultrasonic bath and dried at 80°C for 48h. They were then observed under

Sample Localisation Coordinates

Nature of the sample d of the clast measured (g.cm -3 ) (density range of all clasts)

Vesicularity of the clast (%) (vesicularity range of all clasts)

Bezy 3e

Blast deposits (layer B), cryptodome N55°56.465', E160°36.887' Andesite 1.73 (1.71-2.09) 30.7 (16.6-31.7)

Bezy 8f

Blast deposits, (layer B) cryptodome N55°55 .021', E160°37.826' Andesite 1.57 (1.49-1.83) 37.3 (26.6-40.4)

Bezy 9

Pumices; Plinian phase N55°54.987', E160°38.843' Andesite 1.12-1.52 (0.90-1.94) 42.6-57.9 (26.7-66.0) a binocular microscope to select the fractions in which the crystals were the more abundant and automorphic. Orthopyroxenes (opx), plagioclase (plag), magnetites (mgt) and amphiboles (amph) were then handpicked under the binocular microscope in the different selected size fractions. They were mounted in epoxy resin and polished up to 0.3 µm to the middle part of the crystals. Opx were oriented with c-axis in a north-south direction as they have been used for intracrystalline diffusion modelling [START_REF] Kilgour | Timescales of magmatic processes at Ruapehu volcano from diffusion chronometry and their comparison to monitoring data[END_REF][START_REF] Ostorero | Time -window into the transcrustal plumbing system dynamics of Dominica (Lesser Antilles)[END_REF][START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF]. Before scanning electron microscope (SEM) or electron microprobe micro-analyzers (EPMA) investigations, selected mounts were all carbon-coated. Thin sections were also made in the blast deposits and pumices from the Plinian phase.

V.3.3. Whole rock

V. Whole rock analyses were performed by ICP-OES at Centre de Recherches Pétrographiques et Géochimiques (CRPG) in Nancy (Carignan et al., 2001).

V.3.3.2. Vesicularities and densities estimation

Vesicularities and densities estimations are used to select samples with different textural characteristics.

For this kind of analysis, each clast sample is sawn to obtain a 1 cm wide clast. A first weighing of the part of these 1 cm wide clasts is realized in the air with a precision balance. The principle of buoyancy is used here to know the volume of water displaced by the sample. However, the presence of interconnected vesicles in the samples (vesicularity) constitutes a bias as water penetrates in the pores, but, to avoid this, the samples are immersed in a paraffin bath to avoid the penetration of water into the vesicles (~1h, paraffin heated up to 100 °C). They are then weighed again in the air and then in water.

The density of the fragments is deduced using the following equation (V.1):

𝜌 = 𝑚 𝑎𝑖𝑟 𝑚 𝑝𝑎𝑟𝑎𝑓𝑓𝑖𝑛𝑒𝑑 -𝑚 𝑤𝑎𝑡𝑒𝑟 (V. 1)
where 𝑚 𝑎𝑖𝑟 is the mass of the clast in the air (g), 𝑚 𝑝𝑎𝑟𝑎𝑓𝑓𝑖𝑛𝑒𝑑 , the mass of the clast in the air after the paraffin bath (g) and 𝑚 𝑤𝑎𝑡𝑒𝑟 , the mass of the fragment with paraffin in the water (g).

The vesicularity (in %) of the fragments is found using the equation (V.2):

𝑣𝑒𝑠 = 100 𝐷𝑅𝐸 -𝜌 𝐷𝑅𝐸 (V. 2)
where ρ is the density of the fragment and DRE, "Dense Rock Equivalent", the powder density obtained with a pycnometer.

V.3.4. Crystals textural observations: Scanning Electron Microscope

All crystals were observed under Scanning Electron Microscopes (SEM): Zeiss Supra 55VP (Sorbonne Université, Institut des Sciences de la Terre de Paris, ISTeP, Paris) and the Carl Zeiss EVO MA10 SEM at the PARI platform at the Institut de physique du globe de Paris (Université Paris Cité) using an acceleration voltage of 20kV and a beam current of 8 nA. Opx zoned crystals were identified as they present on back-scattered electrons (BSE) images, grey levels of different intensities. Proportions of zoned/unzoned opx crystals have been determined on all the crystals mounted by looking closely at the SEM images. Higher resolution images (mag x755) were taken for the diffusion modelling and intercalibration with chemical profiles of opx with the same parameters as mentioned above, a high integration line (N=7, 7 integrations of each image line to reduce the impact of noise) and a dwell time per pixel of 48 µs. The melt inclusions in the crystals were also identified by SEM images of sufficient size for electron microprobe analysis (> 10 µm) as well as matrix glass areas larger than 10 µm.

An identification of possible chemical zonations in the mgt (presence of a melt inclusions ring) was also possible in order to position the rim-core profiles to be carried out with the electron microprobe for Fe-Ti oxides.

V.3.5. Compositional analysis of crystals and melt inclusions by EPMA

V. 3.5.1. Point

measurements in crystals

Opx crystals have been analysed for major elements (Si, Al, Ca, Mg, Na, K, Ti, Fe, Mn et P) with an acceleration voltage of 15 kV, a beam current of 10 nA and a focused beam of 2 µm (CAMECA SX-Five and SX-100; Service Camparis, Paris). Counting time on peak and background for Fe and Mg was set at 80s and at 10s for the other elements. The core to rim compositional profiles in zoned opx had a 2 µm step and an average length of 100 µm (~4 hours per profile). They were acquired perpendicular to the long axes of the opx and away from the corners to avoid three-dimensional effects such as growth (Allan et al., 2013;Fabbro et al., 2018;Krimer and Costa, 2017;[START_REF] Ostorero | Time -window into the transcrustal plumbing system dynamics of Dominica (Lesser Antilles)[END_REF]. Four points were measured in the unzoned crystals.

For compositional analyses in magnetite/ilmenite pairs, the counting time was set at 30 s for Fe, Ti and Al and 10 s for the other elements. For mgt profiles, the same analytical conditions as opx were used (15 kV, 10 nA and a focused beam of 2 µm) and a counting time of 10 s was set for all elements except for Fe, Ti and Al (60 s).

V. 3.5.2

. Point measurements in melt inclusions

Using the methodology developed by Balcone-Boissard et al. ( 2018), for the major and minor element measurements of the residual glasses and glassy melt inclusions (MI) of the samples, an acceleration voltage of 15kV was used, with counting times varying from 5 to 30 s depending on the element (Na, Si: 5 s; Cl: 30 s and others: 10 s) and a beam size of 6 µm to avoid alkali migration. To avoid Na loss, a beam current between 6-10 nA was used and Na and Si elements were counted first and simultaneously. All precautions required for MI investigations were applied by selecting the roundshaped ones, suggesting the absence of crystallization after entrapment in the crystals, and without bubbles (possibly indicating outgassing) (Balcone-Boissard et al., 2018;[START_REF] Rose-Koga | Silicate melt inclusions in the new millennium: A review of recommended practices for preparation, analysis, and data presentation[END_REF]. The host mineral compositions have been measured close to the analysed melt inclusions. Internal standards used for glass measurements intercalibration EPMA sessions were analysed in the same conditions as the residual glasses (Supplementary Data V.1). Corrections for the compositional effects of postentrapment modifications were carried out after EPMA analyses (Putirka, 2008).

V.3.6. Temperatures

The WingMIgob software (http://www.jgeosci.org/detail/jgeosci.319) [START_REF] Yavuz | WinMIgob: A Windows program for magnetite-ilmenite geothermometer and oxygen barometer[END_REF] was used to estimate the temperatures, using the geothermometers of titanomagnetite-ilmenite of Sauerzapf et al. (2008) and [START_REF] Ghiorso | Thermodynamics of Rhombohedral Oxide Solid Solutions and a Revision of the FE-TI Two-Oxide Geothermometer and Oxygen-Barometer[END_REF], after passing the test of [START_REF] Bacon | Mg/Mn partitioning as a test for equilibrium between coexisting Fe-Ti oxide[END_REF] (Supplementary Note). This test is based on the Mg/Mn magnetite-ilmenite partitioning, which is valid for fresh volcanic rocks but unvalid for altered volcanic rocks, for example [START_REF] Bacon | Mg/Mn partitioning as a test for equilibrium between coexisting Fe-Ti oxide[END_REF][START_REF] Yavuz | WinMIgob: A Windows program for magnetite-ilmenite geothermometer and oxygen barometer[END_REF] (Supplementary Data V.2).

V.3.7. Timescales of magmatic processes

Interdiffusion timescales were modelled in zoned crystals, which have different rims of different compositions having recorded magmatic processes, as these compositional gradients result from crystallization (causing melt differentiation and crystal zonations) or other processes such as magma mixing [START_REF] Chakraborty | Diffusion chronometry of volcanic rocks: looking backward and forward[END_REF]Costa et al., 2020;Petrone and Mangler, 2021).

An En increase in opx (reverse zonations) is linked to an increase in temperature, with, or without, mixing of the magma with hotter, more basic magma (Martel et al., 1999;[START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF] or to degassing processes with a lower water content in the melt (Waters and Lange, 2017). On the other hand, a decrease in En content (normal zonations) may be related to a drop in pressure, mixing with cooler magma [START_REF] Kahl | Compositionally zoned crystals and real-time degassing data reveal changes in magma transfer dynamics during the 2006 summit eruptive episodes of Mt. Etna[END_REF]Martel et al., 1999;Saunders et al., 2012;[START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF] or a high water content in the melt (Waters and Lange, 2017).

The zonations in crystals can be linked to magmatic environments (ME; compositional plateaus associated to stable storage conditions of the magmas). If the same compositional plateau is found in a significant amount of crystals, it can be concluded that these plateaus were produced by significant changes in the magmatic conditions and can be considered as ME (Kahl et al., 2011;[START_REF] Ubide | Deep Magma Storage Revealed by Multi-Method Elemental Mapping of Clinopyroxene Megacrysts at Stromboli Volcano[END_REF]. The timescales modelled on these ME changes allow to constrain the magma dynamics in time.

The timescales of these magmatic processes can be estimated by using second Fick's law, on compositional profiles, by knowing the initial conditions, boundary conditions and diffusion coefficient of the elements of interest. The diffusion coefficient depends on several parameters, such as the chemical composition of the crystals (Xi; molar fraction of the mineral constituent element), temperature (T in K), pressure (P in Pa), oxygen fugacity (fO2) and water fugacity (fH2O) (Costa et al., 2020;Costa and Morgan, 2010).

V.3.7.1. Fe-Mg interdiffusion modelling in opx

For opx, the same method as in Allan et al. (2013) and Couperthwaite et al. (2020) was followed to model the timescales associated to the zonations in the zoned crystals, using the parametrization of the interdiffusion coefficient D of Fe and Mg elements in opx (Ganguly and Tazzoli, 1994). The Fe-Mg interdiffusion profiles were modelled in one dimension, across the c-axis of the opx , parallel to the baxis of the crystals (Allan et al., 2013;Couperthwaite et al., 2020). D has been defined without an oxygen fugacity fO2 dependence in equation ( 3) (Ganguly and Tazzoli, 1994) by the equation (V. The fO2 dependence on Fe-Mg interdiffusion coefficient is still under discussion. Ganguly and Tazzoli (1994) did not constrain the fO2 oxygen dependence but they hypothesized it by analogy to the behavior of olivine (exponent of 1/6). The fO2 correction was added in Allan et al. (2013) A series of experimental determinations of Fe-Mg interdiffusion in En-rich opx (En91 and En100) constrained a lower fO2 dependence than inferred from Ganguly and Tazzoli (1994) (with an exponent between 0 to 0.05), but that depended on the En composition of the crystal (Dohmen et al., 2016). Using their coefficient determined for highly magnesian compositions would then lead to effects of composition on diffusivity, yielding erroneous timescales (Allan et al., 2017). As Allan et al. ( 2017), we thus used the coefficient of Ganguly and Tazzoli (1994), that includes the compositional range of the opx of this study (En59-68; Supplementary Figure V.3), in the absence of more recent determination close to the conditions of our opx and without the oxygen dependence (as it was shown to be weak and not constrained for a wide range of opx compositions (Dohmen et al., 2016)).

However, to study the effect of fO2 and to see the differences in the diffusion timescales by using another diffusion coefficient, we modelled the timescales first without considering fO2 and then discussed about its effect (with The determination of the temperature, the diffusion coefficient and its measurement are the main sources of uncertainties (Costa and Morgan, 2010;Morgan et al., 2004). Uncertainties on the calibration of measurements and point spacing during EPMA analyses, as well as the resolution of SEM images are also sources of uncertainties [START_REF] Ostorero | Time -window into the transcrustal plumbing system dynamics of Dominica (Lesser Antilles)[END_REF]. Overall, the measurement of the diffusion coefficient and temperature represent the main sources of uncertainty (Costa and Morgan, 2010;Solaro-Müller, 2017;[START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF].

To compare all the timescales depending on the parameters, we evaluated each dataset via a Gamma distribution (Poisson processes), as eruptive processes and magmatic events occur at random times and given the random nature of crystal sampling [START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF]. The distribution of each population of timescales is studied, to determine the expected timescale depending on the parameters chosen, i.e. the probability of a random magmatic process to occur, calculated from the probability distribution. The standard error estimates are also obtained (see Appendix). A goodness-of-fit test was also made to check that the Gamma distribution represents the data, returning a probability (p-value) at a given confidence of 0.05 (5%, p-values greater than 0.05 passed the test. All the details on the functions used are detailed in [START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF].

V. 3.7.2

. Ti modelling in mgt

Fe-Ti oxides, and mgt in particular, are known to rapidly re-equilibrate (within a few days to months) in response to fluctuations in magmatic conditions such as temperature, surrounding fluid composition or oxygen fugacity fO2 influencing the partitioning of elements at equilibrium (Allan et al., 2017;[START_REF] Aragon | Cation diffusion in titanomagnetites[END_REF][START_REF] Coombs | Reaction rim growth on olivine in silicic melts: Implications for magma mixing[END_REF][START_REF] Devine | Magma Storage Region Processes Inferred from Geochemistry of Fe-Ti Oxides in Andesitic Magma, Soufriere Hills Volcano, Montserrat[END_REF][START_REF] Freer | An experimental study of magnetite-titanomagnetite interdiffusion[END_REF]Martel et al., 1999;Nakamura, 1995a;[START_REF] Sievwright | Diffusion and partition coefficients of minor and trace elements in magnetite as a function of oxygen fugacity at 1150 o C. Contrib[END_REF][START_REF] Venezky | Petrology and Fe-Ti oxide reequilibration of the 1991 Mount Unzen mixed magma[END_REF].

Ti diffusion is studied in mgt by using the diffusion coefficient relationship by [START_REF] Aragon | Cation diffusion in titanomagnetites[END_REF], which is described in the equation (V.5) below:

𝐷 𝑇𝑖 * = 𝐷 0 0 𝑒 -( 𝐸 𝑎 𝑅𝑇 ) + 𝐷 𝑉 0 𝑒 -( 𝐸 𝑣 𝑅𝑇 ) 𝑓 𝑂 2 2/3 -𝐷 𝐼 0 𝑒 -( 𝐸 𝐼 𝑅𝑇 ) 𝑓 𝑂 2 -2/3 (V. 5)
For T of interest and fO2 of interest and where 𝐷 0 0 = 8.2.10 -3 cm 2 /s ; 𝐷 𝑣 0 = 1. 75.10 -12 (cm 2 atm 2/3 )/s; 𝐷 𝐼 0 = 3.9.10 6 (cm 2 atm 2/3 )/s; 𝐸 0 = 60 Kcal/mol; 𝐸 𝑣 = -28 Kcal/mol and 𝐸 𝐼 = 159 Kcal/mol. The first member of this expression corresponds to the migration of Ti into the crystal structure while the second and third describe the rate at which new sites are available for the Ti cation [START_REF] Aragon | Cation diffusion in titanomagnetites[END_REF].

If magnetites are unexsolved, Ti content may be used to constrain T and/or fO2 changes during the late stages of magma ascent in the conduit and magnetite crystallization, shortly before quenching (timescales of the order of days to months) (Boudon et al., 2015;Nakamura, 1995a). Ti profiles of unexsolved magnetites were obtained using first SEM images on magnetites hand-picked in the 315-500 and 250-315 µm fractions to identify the best profiles for diffusion modelling. Ti profiles are not visible on SEM images by variations of grayscale intensities as in zoned opx (low contrast in composition) but the occurrence of a ring of melt inclusions in magnetites allows the identification of suitable magnetites (Boudon et al., 2015) (Supplementary Figure V.5). Then, quantitative Ti profiles were analysed after EPMA sessions. Diffusion profiles are between 100 and 180 µm long, from core to rim with a 2 µm step. Then Ti diffusion profile may be fit by diffusion models at 915 ± 17°C for dacites and 991 ± 44 °C for andesites and fO2 of ΔNNO + 1.6 ± 0.4 for the blast and fO2 of ΔNNO + 1.8 ± 0.3 for the Plinian phase (fO2 of 5.5.10 -10 Pa and 1. 1.10 -7 Pa, respectively) (Sauerzapf et al., 2008) (Table V.1), with the diffusion coefficient depending mainly on the temperature, fO2 and Ti composition [START_REF] Aragon | Cation diffusion in titanomagnetites[END_REF]Sauerzapf et al., 2008). The initial conditions are a step function, using a spreadsheet with a diffusion model of D. J. Morgan (unpublished). Uncertainty is estimated by propagating the temperature uncertainties onto the timescale.

V.4. Results

In this study, we used samples from the climactic phase of the eruption of 1955-1956 of Bezymianny (from the blast and from the Plinian phase).

V.4.1. Vesicularities and densities estimation

Densities and vesicularities of the clasts from the blast and the Plinian phase studied here have been measured (Table V.1). For the blast, clast densities vary between 1.5-2.1 kg.m -3 (vesicularity between 20-40 %) and the two clasts studied here have a density between 1.6-1.7 kg.m -3 (vesicularity between 30-40%). The pumices from the Plinian phase have a density between 1-2 kg.m -3 (vesicularity between 30-70 %) and the samples studied here have a density of 1.1-1.5 kg.m -3 (vesicularity between 43-58 %) (Table V.1).

V.4.2. Whole rocks

Whole rock compositions from the clasts of the blast and Plinian phase are andesites, with ~60 wt% SiO2 and 5 wt% Na2O +K2O (recalculated with total iron as FeO and on anhydrous basis) (Figure V.2; Supplementary Data V.1).

V.4.3. Mineral assemblage

The mineral assemblage of the clasts from the blast and the Plinian phase is identical and consist of plag, amph, opx and mgt (in decreasing order of proportion). For the blast, plag and amph compositions were not investigated but for the Plinian phase, plag anorthite contents (𝐴𝑛 = 𝐶𝑎 𝐶𝑎+𝑁𝑎+𝐾 ) are between An40-75, amph are magnesio-hornblende, and opx and mgt are presented in more details below (Supplementary Data V.1).

V.4.4. Glass: residual glasses and melt inclusions (MI)

Blast MI trapped in cores of opx of the blast have a rhyolitic composition (72-77 wt% SiO2, 5-6 wt% Na2O+K2O; Figure V.2; Supplementary Figure V.6; Supplementary Data V.1). The residual glasses also plot in the rhyolitic domain (72.8-80.8 wt% SiO2, 6.9-7.6 wt% Na2O+K2O; Figure V.2; Supplementary Data V.1). In the K2O, Al2O3, CaO and FeO versus SiO2 diagrams, the compositions of the MI and residual glasses are in the same domain (Supplementary Figure V.7). No significant difference between MI hosted in opx is present, with the En content of the hosted opx between En63-67 and CaO and Al2O3 contents between 1.9-2.

V.4.5. Temperatures and oxygen fugacity of the magma

Blast

Temperatures have been estimated by using the compositions of magnetite and ilmenite pairs and are of 915 ± 17 °C using Sauerzapf et al. (2008) and 935 ± 19 °C using Ghiorso and Evans (2008) (on 6 pairs; Table V.2; Supplementary Data V.2) [START_REF] Yavuz | WinMIgob: A Windows program for magnetite-ilmenite geothermometer and oxygen barometer[END_REF]. The oxygen fugacity fO2 is of 0.9-1.6 log units above the nickel-nickel oxide (NNO) buffer (Table V.2; Supplementary Data V.2) (Sauerzapf et al., 2008;[START_REF] Yavuz | WinMIgob: A Windows program for magnetite-ilmenite geothermometer and oxygen barometer[END_REF].

Plinian phase

With the compositions of magnetite and ilmenite pairs, temperatures have been estimated with Sauerzapf et al. (2008) at 991 ± 44 °C or 1020 ± 50 °C using [START_REF] Ghiorso | Thermodynamics of Rhombohedral Oxide Solid Solutions and a Revision of the FE-TI Two-Oxide Geothermometer and Oxygen-Barometer[END_REF] (on 20 pairs; Table V.2; Supplementary Data V.2) [START_REF] Yavuz | WinMIgob: A Windows program for magnetite-ilmenite geothermometer and oxygen barometer[END_REF]. The oxygen fugacity fO2 is of 1-1.8 log units above the NNO (Table V.2; Supplementary Data V.2) (Sauerzapf et al., 2008;[START_REF] Yavuz | WinMIgob: A Windows program for magnetite-ilmenite geothermometer and oxygen barometer[END_REF]. The temperatures were also estimated based on the melt inclusions hosted in the opx and are of 930 ± 26 °C (Table V.2; Supplementary Data V.1).

The temperatures used in the timescales modelling are those of the geothermometer of Sauerzapf et al. (2008) presented before for the blast and Plinian phase and those obtained with the other geothermometers (using [START_REF] Ghiorso | Thermodynamics of Rhombohedral Oxide Solid Solutions and a Revision of the FE-TI Two-Oxide Geothermometer and Oxygen-Barometer[END_REF] for mgt/ilm or opx/liq (Putirka, 2008); Table V.2

V.4.6. Opx

V. 4.6.1. Zonations types Based on SEM images, identification of the textures and zonations of the opx was done. Single-zoned opx were identified, with either a normal-zoned rim, with a Mg-rich core (darker zone on a grayscale image) and Fe-rich rims, or reverse-zoned rim (with a Fe-rich core and Mg-rich rims) but also multiplezoned opx (

Sample

Average V. 4.7.3. Timescales mgt Using the temperature estimated with Sauerzapf et al. (2008), Ti diffusion timescales were modelled in the blast and Plinian phase (Figure V.7; Supplementary Data V.5).

Blast

Out of the 637 magnetites, 10 profiles have been measured (Supplementary Data V.5). Timescales modelled on reversely-zoned mgt are young, shorter than 200 days, with 90 % of the timescales from 2-1 months +50/ -5 days (Figure V.7c; Supplementary Data V.5).

Plinian phase

Out of the 644 magnetites, 29 profiles have been measured (Supplementary Data V.5). Timescales of the Plinian phase are shorter than the ones of the blast, with timescales of less than 15 days prior to the climactic phase, with the majority of them ranging from 6-2 days +11/ -1 days (Figure V.7d; Supplementary Data V.5).

Only two timescales were modelled on the core to the inner reverse-zoned rim of multiple-zoned magnetites in the blast and Plinian phase; they give similar timescales (Supplementary Figure V.13).

V.5. Discussion

Magma dynamics are important processes to decipher prior to an eruption, to better understand the unrest at a specific volcano. We are interested in the climactic phase of the 1955-1956 eruption of Bezymianny, which is charcaterised by the laterally directed blast and the Plinian phase that occurred on 30 March 1956. The preclimactic phase (eruptions of Vulcanian-type) in October-November 1955 is not studied as the magmas might be unrelated to the ones involved after the flank collapse, as evidenced by the different crystal assemblages from the different deposits and rims disequilibria of amph in cryptodome bombs [START_REF] Gorshkov | Gigantic eruption of the volcano bezymianny[END_REF][START_REF] Gorshkov | Bezymianny Volcano and Peculiarities of Its Last Eruptions in[END_REF][START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF] (Supplementary Figure V.2). The pre-eruptive processes deciphered here correspond to the period recorded by opx and mgt up to the flank collapse and explosive activity, combined with the petrological information recorded by plag and amph from other studies (d'Augustin, 2021;[START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF][START_REF] Shcherbakov | Phase equilibria constraints on preeruptive magma storage conditions for the 1956 eruption of Bezymianny Volcano, Kamchatka, Russia[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF].

V.5.1. The crustal reservoir of the 1956 eruption

The melt inclusions data allow to refine the initial state of the plumbing system and the origins of the magma of the blast and Plinian phase. 26 melt inclusions were analysed in opx and other crystals from the Plinian phase (d'Augustin, 2021), they can give information on the pressure conditions of the magmatic storage (Supplementary Figure V.14; Supplementary Figure Data V.1). Using their CO2 and H2O content, in an isobaric solubility plot for the temperature estimated in the MI, pressure ranges are roughly estimated to be between 250-400 MPa (approximately between 9-15 km depth, for a density of 2700 kg.m -3 ) (Supplementary Figure V.14) [START_REF] Newman | VolatileCalc: a silicate melt-H2O-CO2 solution model written in Visual Basic for excel[END_REF]. According to the plumbing system models and seismic data, two magma chambers are present at different depths (between 8-11 km (200-300 MPa) and 19-26 km (500-750 MPa)) (Almeev et al., 2013a(Almeev et al., , 2013b[START_REF] Almeev | Problems of the Stoichiometry and Thermobarometry of Magmatic Amphiboles: An Example of Hornblende from the Andesites of Bezymyannyi Volcano, Eastern Kamchatka[END_REF][START_REF] Fedotov | The magmatic system of the Klyuchevskaya group of volcanoes inferred from data on its eruptions, earthquakes, deformation, and deep structure[END_REF][START_REF] Kadik | Physicochemical Crystallization Conditions and Genesis of Andesites[END_REF][START_REF] Koulakov | Three different types of plumbing system beneath the neighboring active volcanoes of Tolbachik, Bezymianny, and Klyuchevskoy in Kamchatka[END_REF][START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF][START_REF] Ter Heege | Experimental Determination of Fe-Mg Interdiffusion Coefficients in Orthopyroxene Using Pulsed Laser Ablation and Nanoscale Thin Films[END_REF][START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]. Following the melt inclusions data, the magma that fed the Plinian phase may have been located in the upper magma chamber. For the MI in the blast, no CO2 or H2O were measured as only 10 melt inclusions were found in the opx so no estimates of the pressures/depth can be assessed but we can infer a more or less same depth as the Plinian phase, as the core compositions of crystals between the two deposits are the same (Figure V.5).

Furthermore, [START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF] concluded that the magmatic chamber that fed the 30 March 1956 climactic phase endured significant changes that led to a modification of the phenocryst assemblage with a fresh magma injection that caused the significant heating of the magmatic chamber, amphibole reaction rims less than 37 days before the blast and eventually, the catastrophic eruption [START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF]. So, the magma that led to the blast may have been stored in the shallower reservoir. [START_REF] Shcherbakov | Phase equilibria constraints on preeruptive magma storage conditions for the 1956 eruption of Bezymianny Volcano, Kamchatka, Russia[END_REF] showed, following amphibole breakdown rims in pumiceous pyroclastic deposits from the Plinian phase, that the magma was transiently stored at ~3 km depth prior to the final ascent for at least 40 days. But we do not believe that the magma at the origin of the Plinian phase was stored as shallowly as 3 km as, otherwise, the Plinian phase would not have occurred so, hereafter, the shallow reservoir refers to the depth between 9-15 km depth inferred from the melt inclusions.

V.5.2. Magmatic dynamics and timescales

As seen before, all magma that erupted during the eruption was initially stored in the same shallow reservoir (~9-15 km depth). Opx and mgt zonations give us insights into the magmatic processes experienced by the magma, either stored in the shallow reservoir and then in the cryptodome (for the magma that led to the blast) or still in the shallow reservoir (magma emitted during the Plinian phase).

Approximately 80 % of the opx are zoned, indicating that they experienced significant magmatic changes before the climactic phase. The small proportion of unzoned opx, for both the blast and Plinian phase (20 %; Figure V.4) are related to crystals that were in areas in the magmatic reservoir that did not record the magmatic processes experienced by the zoned opx, as the unzoned opx and the cores of the zoned ones have the same compositions (Supplementary Figure V.10). Some MZ crystals recorded zonations in the same ME (one crystal for the blast and three for the Plinian phase), interpreted as magma movement in the same reservoir (Supplementary Data V.4).

Blast

MZ opx are dominant in the blast, with a majority of normal-and then reverse-zoned opx for the last rims (Figure This timescale coincides with the flank uplift and deformation from the end of November 1955 [START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF][START_REF] Gorshkov | Gigantic eruption of the volcano bezymianny[END_REF]. Then, in the last month prior to the eruption up to 3 months before the climactic phase, heating from below or degassing of the magma within the cryptodome led to the development of reverse zoning in the last rims of the opx (Waters and As we inferred that the magma that gave the blast and the Plinian phase was initially stored in the same shallow reservoir (~9-15 km), the reactivation of the reservoir may also have occurred a few years before the climactic phase, as for the blast. This reactivation was recorded by few crystals from the blast and no crystals from the Plinian phase, maybe because of the sampling bias when studying crystals.

The timescales modelled in these zoned opx show magmatic processes up to 6 months prior eruption, with normal or reverse zoning evidenced in inner rims (Figure V.6c). As for the blast, the inner rims correspond to self-mixing with convection in the magmatic reservoir (Couch et al., 2001). The N-1 and N rims also recorded normal zoning, with the last N rims having timescales less than 3 months, mainly below 1 month (Figure Magma leading to the Plinian phase thus recorded the magmatic activity in the shallow reservoir, with part of the magma that began to be affected by decompression, while its base was fed by hotter magma from the deeper reservoir. When the flank collapse occurred, it was followed by the laterally directed blast that led to a strong decompression event in the volcanic conduit. This decompression generated a fast magma ascent of the magma that gives the Plinian phase.

V.5.3. Final pre-eruptive scenario

We propose the following pre-eruptive scenario prior to the climactic phase, by considering all the crystal assemblage of the deposits (opx and mgt from this study and plag and amph from other studies), as well as the storage conditions deciphered by the melt inclusions and residual glasses (Figure V.8).

1) 3 years to 6 months prior the climactic phase (30 March 1956), magma in the upper reservoir of Bezymianny began to self-mix, forming the MZ opx: decompression of the upper part of the reservoir forms normal zonations and heating of the base of the reservoir forms reversed zoned rims in opx (Figure V.9a). 2) Then, a seismic crisis occurs from September to October 1955, with phreatomagmatic activity with Vulcanian-type ash emissions. The magma involved in this preclimactic phase is different than the magma that formed the cryptodome, as their whole rock compositions and crystal assemblages are different [START_REF] Gorshkov | Gigantic eruption of the volcano bezymianny[END_REF][START_REF] Gorshkov | Bezymianny Volcano and Peculiarities of Its Last Eruptions in[END_REF][START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF]. A small pocket of magma stored above the shallow reservoir may have been activated to produce this preclimactic phase (Figure V.2; 9). On the other hand, in the shallow reservoir where the magma at the origin of the blast and the Plinian phase is initially stored, normal-zoned rims in the opx of the blast are formed, that can correspond to the beginning of the ascent of the magma, leading to some decompression in the reservoir (Figure V.9b). 3) 6 to 5 months before the climactic phase, from the end of November to January, magma at the origin of the blast is stored in the cryptodome (Figure V.8; Figure V.9c). This magma may have been stored within the flank of the volcano at a depth of about 1 km away from the summit (as the height reduction after the blast was of ~800 m) [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF][START_REF] Tokarev | The prediction of large explosions of andesitic volcanoes[END_REF]. This magma began to cool down, as it was in contact with colder material near the summit, creating normal zonations in the opx (Figure V.9c). This cryptodome formation is accompanied by flank deformation and uplift, as well as lava dome growth in the summit crater [START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF][START_REF] Gorshkov | Gigantic eruption of the volcano bezymianny[END_REF][START_REF] Gorshkov | Bezymianny Volcano and Peculiarities of Its Last Eruptions in[END_REF]. Rare and weak ash emissions recorded from November to March occurred during lava dome growth [START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF][START_REF] Gorshkov | Gigantic eruption of the volcano bezymianny[END_REF][START_REF] Gorshkov | Bezymianny Volcano and Peculiarities of Its Last Eruptions in[END_REF]. 4) One to 2 months before the climactic phase, around January-February, magma degassing or heating processes occurred in the cryptodome, forming the breakout rims of the amph (granular or symplectitic textures as well as volumetric decomposition (Supplementary Figure V This could correspond to a magma influx from the deeper reservoir evidenced from petrological and geophysical data (from 19 to 26 km), pushing magma from the deeper one to the shallower reservoir where the magma giving the Plinian phase is stored (Figure V.9e). This magma influx could have been triggered by a mafic magma injection from even deeper areas (below 30 km). The injected mafic magma did not interact with the magma in the shallower reservoir, as no mafic enclaves were found in the 1956 climactic phase deposits [START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF]. Furthermore, no major En changes were produced in the opx and their last rims compositions dominantly return to the composition of their cores, showing that no mafic magma was sampled and that the magma influx only participated in the heating of the shallower reservoir (Figure 7) The lavas from the lava dome of August 1956 are still andesites [START_REF] Gorshkov | Gigantic eruption of the volcano bezymianny[END_REF], meaning that all the andesitic magma stored in the shallower reservoir did not erupt in March 1956. After 1956, the more mafic magma at the base of the shallow reservoir, fed from the deep reservoir, participates in emitting less evolved magmas during the following eruptions [START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF].

It is difficult to compare the magma dynamics with the seismicity as it is probable that a large part of the precursors were not detected, respect to the unique seismic station that was far away from Bezymianny (at 42 km from the volcano) [START_REF] Tokarev | The prediction of large explosions of andesitic volcanoes[END_REF][START_REF] Tokarev | Volcanic Earthquakes of Kamchatka[END_REF]. The pre-eruptive seismic crisis from September to October 1955 likely corresponded to magma ascent giving the first Vulcanian explosions. So, it is probable that the seismic crisis began before, perhaps 3 years before the climactic phase, in relation with the first timescale recorded in the opx.

V.5.4. Comparison with other directed blast eruptions

Timescales for the magma of the cryptodome accumulation identified here thus begin around 6 to 5 months prior to the climactic phase. They correspond to the beginning of the observed deformation of the flank of the volcano (6 months before) by the injection of the magma in the cryptodome [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF][START_REF] Belousov | Bezymyannyi Eruption on March 30, 1956 (Kamchatka): Sequence of Events and Debris-Avalanche Deposits[END_REF]. We can compare Bezymianny 1956 climactic phase with other eruptions that generated laterally directed blasts following the injection of a magma in a cryptodome. Two eruptions of this type occurred in the last decades (1980 eruption of Mount St. Helens, USA and 1951 eruption of Mount Lamington, Papua New Guinea). For the eruption of Mount St. Helens which was of the same magnitude as 1956 Bezymianny, devastating areas of 600 km 2 (500 km 2 for Bezymianny), pre-eruptive timescales prior the flank collapse and the laterally directed blast were constrained by the opx timescales [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF][START_REF] Lipman | The 1980 eruptions of Mount St[END_REF][START_REF] Saunders | [END_REF]. For Mount St. Helens, seismic swarms mainly occurred 2 months before the flank collapse [START_REF] Endo | Locations, magnitudes and statistics of the March 20-May 18 earthquake sequence[END_REF]Saunders et al., 2012;[START_REF] Scandone | Magma supply, magma ascent and the style of volcanic eruptions[END_REF] and the observation of the deformation began 50 days before [START_REF] Christiansen | Chronology of the 1980 eruptive activity[END_REF][START_REF] Moore | Topographic and srictural changes, MArch-July 1980photogrammetric data[END_REF]. Magma dynamics inferred for the 18 May 1980 eruption of Mount St. Helens is close to what is inferred here for 1956 Bezymianny, with recharge or convective overturn of the magmatic system approximately one year before the eruption (Saunders et al., 2012). For Mount Lamington, the situation was similar to Bezymianny with perhaps the growth of a small lava dome and/or the injection of a part of the magma in the northern flank of the volcano [START_REF] Belousov | The 1951 eruption of Mount Lamington, Papua New Guinea: Devastating directed blast triggered by small-scale edifice failure[END_REF][START_REF] Lacroix | La Montagne Pelée et ses éruptions[END_REF][START_REF] Taylor | The 1951 eruption of Mount Lamington, Papua[END_REF]. The unstable flank of the volcano collapsed rapidly and a violent explosion with a vertical component occurred on 21 January 1951, generating a laterally directed blast. As no monitoring existed on this volcano, the only indication is the beginning of seismic tremors felt by the population at the beginning of the year [START_REF] Belousov | The 1951 eruption of Mount Lamington, Papua New Guinea: Devastating directed blast triggered by small-scale edifice failure[END_REF]. No timescales modelling was carried out in crystals. A smaller blast occurred on Soufriere Hills volcano (Montserrat) in December 1997, the situation is different as it is the voluminous growing lava dome that exerted a load on the hydrothermalized and weathered southern flank. The collapse of this flank of the volcano generated a depressurization of the inner part of the lava dome that produced a laterally directed blast of one order of magnitude smaller compared to 1956 Bezymianny and 1980 Mount St. Helens [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF][START_REF] Hart | Discharge of pyroclastic flows into the sea during the 1996-1998 eruptions of the Soufriere Hills volcano[END_REF][START_REF] Sparks | Generation of a debris avalanche and violent pyroclastic density current on 26 December (Boxing Day) 1997 at Soufrière Hills Volcano[END_REF]. But this event occurred 2 years after the growth of the first lava dome.

Pre-eruptive magma dynamics studies on the following open-system eruptive activity of Bezymianny, would be interesting to see if a systematic pattern of the timescales reactivation can be identified (same spatio-temporal mgma dynamics identified before several eruptions), as it was shown for several recent eruptions of this volcano from 2006-2012 eruptions (Davydova et al., 2018).

Directed blasts are then threatening and understanding their pre-eruptive dynamics is of particular importance, all the more if little unrest activity is detected prior to these events.

V.6. Conclusion

Combining information from several petrological clocks provide more insights in the pre-eruptive dynamics of catastrophic eruptions. Cryptodome storage occurred around 5-6 months prior to the climactic phase of 30 March 1956 of Bezymianny, followed by a heating process or degassing, mostly during the last two months prior to the flank collapse. Magma at the origin of the Plinian phase also recorded this degassing and/or heating event very shortly prior eruption (in the less 10 days). Here, we have thus estimated the time between the onset of the flank deformation by cryptodome growth and the climactic phase of the eruption; this timescale of less than 5 months has to be considered in risk mitigation. Unrest phases at gravitational instable volcanoes with lava domes are thus to be closely monitored, as lateral blasts can possibly occur.

V.7. Acknowledgments

We would like to thank S. Hidalgo for her help with sample preparation, O. Boudouma and S. Borensztajn for SEM imaging and E. Delairis for the realization of the thin sections, as well as M. Fialin and N. Rividi for assistance during EPMA analyses. We thank L. Loïodice for participating in the study of opx crystals during her thesis of Master 1. We also thank D. J. Morgan for providing the files to model the timescales of the opx and D. Jessop for his help regarding the modelling of the timescales using another diffusion coefficient and for adapting the code from [START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF]. This project was funded by the Institut de physique du globe de Paris (doctoral grant from the French Ministry of Higher Education and Research and Innovation) and ANR V-Care-18-CE03-0010.

V.8. Author contributions

V.9. Additional information

Competing interests: All the authors declare no competing interests.

V.10. Appendix: parametric study of the timescales

Critical parameters for deciphering the timescales in diffusion profiles are the temperature T and oxygen fugacity fO2, as changes of these parameters have effects on the timescales modelled (Supplementary Data V.4).

For the blast and Plinian phase, a higher temperature of 935 or 1020 °C [START_REF] Ghiorso | Thermodynamics of Rhombohedral Oxide Solid Solutions and a Revision of the FE-TI Two-Oxide Geothermometer and Oxygen-Barometer[END_REF] used in the diffusion model with the diffusion coefficient expression of Ganguly and Tazzoli (1994) leads to younger timescales (for the blast: maximum of 790 days compared to 1130 days using 915 °C (Sauerzapf et al., 2008); Table V.2; Supplementary Data V.4). A lower temperature (930 °C, for the Plinian phase) (d'Augustin, 2021) leads to longer timescales (maximum of 1020 days compared to 570 days using Sauerzapf et al. ( 2008)) (Supplementary Figure V.16).

If we consider the fO2, the timescales found are radically younger (using the temperature estimated with Sauerzapf et al. (2008)). For the blast, 80 % of the timescales are below 15 days with the fO2 whereas they are below 140 days without considering the fO2 (Supplementary Figure V.16). For the Plinian phase, they are below 8 days compared to 75 days (without the fO2) (Supplementary Figure V.16). The timescales are longer by at least a factor 5 if we consider the diffusion coefficient of Dohmen et al. (2016) (with the fO2) compared to Ganguly and Tazzoli (1994) diffusion coefficient, with the fO2 (with T of Sauerzapf et al. (2008)) (Supplementary Figure V.16). Compared to Ganguly and Tazzoli (1994) without the fO2, the difference is less visible but as there are two variables: the diffusion coefficient and the presence or not of the fO2 dependence, it is difficult to compare them (Supplementary Figure

V.16).

Using [START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF] method investigating the timescales distributions, expected timescales of these distributions are found (the populations of the timescales all passed the goodness-of-fit test (pvalues > 0.05) (Supplementary Figure V.17-18) [START_REF] Gibbons | P-Values: Interpretation and Methodology[END_REF]).

Comparing the expected timescales of the opx obtained by testing the different parameters, to the median of the timescales, they are longer than the median (84 days for the blast compared to 52 days) and 26 days, for the Plinian phase (± 0.3 days) compared to 13 days, respectively, using the T estimated by Sauerzapf et al. ( 2008)) (Appendix Table V.1; Supplementary Figure V.17). Expected timescales modelled with a higher temperature are 2 times younger for both the blast and Plinian phase, and 3 times longer, using a lower temperature, for the Plinian phase (Appendix Table V.1; Supplementary Figure V.17). Considering the fO2 leads to timescales that are 9-10 times younger than without considering it, for the blast and Plinian phase, respectively (Appendix Table V.1; Supplementary Figure V.17). Changing the diffusion coefficient leads to an expected timescale that is 7 times longer compared to D of Ganguly and Tazzoli (1994) with fO2, for the blast (Appendix Table V 2008) will be discussed in details, so that the same method to estimate the temperature is used.

So, the temperature and oxygen fugacity have a critical influence in the timescales modelling, as well as the diffusion coefficient parametrization. We thus decided to only discuss the timescales modelled with the diffusion coefficient of Ganguly and Tazzoli (1994), that was parametrized with the appropriate En content of the opx as studied here and without the fO2 as its influence was not clearly constrained by the authors. We used the temperatures determined by the same method (in magnetite/ilmenite pairs (Sauerzapf et al., 2008;[START_REF] Yavuz | WinMIgob: A Windows program for magnetite-ilmenite geothermometer and oxygen barometer[END_REF]) (Table V.2), using the temperature which was closer to the other temperature found by other studies in the same samples [START_REF] Almeev | Problems of the Stoichiometry and Thermobarometry of Magmatic Amphiboles: An Example of Hornblende from the Andesites of Bezymyannyi Volcano, Eastern Kamchatka[END_REF][START_REF] Kadik | Physicochemical Crystallization Conditions and Genesis of Andesites[END_REF][START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF][START_REF] Shcherbakov | Phase equilibria constraints on preeruptive magma storage conditions for the 1956 eruption of Bezymianny Volcano, Kamchatka, Russia[END_REF] obtained with the geothermometer of Sauerzapf et al. (2008).

V.A. Supplementary Note

V.A.1. Temperature estimates

The thermo-oxybarometer of titanomagnetite-ilmenite was used to estimate temperatures of equilibration of the magmas (Sauerzapf et al., 2008;[START_REF] Yavuz | WinMIgob: A Windows program for magnetite-ilmenite geothermometer and oxygen barometer[END_REF]. This thermometer is based on the exchange of Fe 2+ + Ti 4+ for 2 Fe 3+ between the solid solutions of titanomagnetite and ilmenite, which is depending on the temperature, following this reaction:

Fe2TiO4 (ulvöspinel) + Fe2O3 (hematite) = FeTiO3 (ilmenite) + Fe3O4 (magnetite).

Regarding the oxybarometer, it is based on iron redox equilibria, involving the Ti-rich endmembers of the two Fe-Ti oxide series: 6 FeTiO3 (ilmenite) + 2 Fe3O4 (magnetite) = 6 Fe2TiO4 (ulvöspinel) + O2 (Sauerzapf et al., 2008).

V.A.2. Timescales modelling

It consists in the intercalibration of chemical composition profiles of zoned opx measured with an electron microprobe, with grayscale profiles on high-resolution north-south oriented images taken with SEM (magnification of 755) (Supplementary Figure V.5). The Mg number is the parameter that seems to best control the grayscale of the images (Couperthwaite et al., 2020;Solaro-Müller, 2017;[START_REF] Petrone | [END_REF]Reed, 2005;Morgan et al., 2004). The ImageJ software (https://imagej.nih.gov/ij/; version 1.52a) is used to produce the grayscale profiles and they are averaged over a certain width to reduce the noise (between 20 to 50 µm, depending on the crystal rim size) [START_REF] Ostorero | Time -window into the transcrustal plumbing system dynamics of Dominica (Lesser Antilles)[END_REF][START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF]. After this intercalibration, the profiles are modelled to obtain the timescales of interdiffusion.

Based on a database of contrasts of particular compositions calculated using a finite difference method, the curve that best fits the one obtained by intercalibration will be searched in this database (Couperthwaite et al., 2020). After rescaling the two curves, the timescale corresponding to the zonation is obtained, as the scaling factor depends on the ratio between the timescales of the database and the time associated with zoning in the crystal [START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF].

The main assumption of the model is that the initial profile between two zones follows a step function, which is modified by diffusion to form sigmoidal concentration gradients and, with time, this steep profile becomes broader, due to interdiffusion of the different elements. During the eruption, the profiles are quenched due to a major drop in temperature. Maximum timescales are then estimated (Morgan et al., 2004).

V.A.3. Equilibrium test for MI in opx

For opx, we used the Rhodes' diagram to test for equilibrium for orthopyroxene, using the following equation (V. A. 1):

𝐾 𝐷 (𝐹𝑒 -𝑀𝑔) 𝑜𝑝𝑥-𝑙𝑖𝑞 = 0.4805 -0.3733 𝑋 𝑆𝑖 𝑙𝑖𝑞 (V. A. 1)
Where 𝑋 𝑆𝑖 𝑙𝑖𝑞 is the cation fraction of SiO2 in the equilibrium liquid. The average KD(Fe-Mg) opx-liq for the inclusions is KD(Fe-Mg) opx-liq = 0.2±0.06, which implies < 1% of recrystallization. The Rhodes diagram in Supplementary Figure V. 6 shows the Mg# of the liquid (measured in the melt inclusion) compared to the Mg# of the orthopyroxene (Putirka, 2008).

V.A.4. ∆ 𝑲 𝑫 𝑭𝒆 𝟐+ -𝑴𝒈 estimations

An experimentally calibrated relationship between the 𝐾 𝐷 𝐹𝑒 2+ -𝑀𝑔 𝑜𝑝𝑥 -𝑙𝑖𝑞𝑢𝑖𝑑 and the H2O content of the melt was found by Waters and Lange (2017) and they showed that an increase of the Mg# of the opx was produced when the content of H2O in the melt decreases, for a series of H2O fluid-saturated phase-equilibrium experiments. Based on their data, the following equation can be derived (V. A. 2):

𝑲 𝑫 𝑭𝒆 𝟐+ -𝑴𝒈 𝑜𝑝𝑥 -𝑙𝑖𝑞𝑢𝑖𝑑 = 0.0979 (𝑤𝑡% 𝐻 2 O) -0.0007 (V. A. 2) The variation between the 𝐾 𝐷 𝐹𝑒 2+ -𝑀𝑔 of the cores and the rims of opx can be calculated in the blast and Plinian phase normal-zoned and reverse-zoned rims, compared to the cores, using the equation (V. A.3), as in Utami et al. (2021):

∆ 𝑲 𝑫 𝑭𝒆 𝟐+ -𝑴𝒈 = ( [𝑋 𝑀𝑔𝑂 𝑙𝑖𝑞𝑢𝑖𝑑 𝑋 𝐹𝑒𝑂 𝑜𝑝𝑥 ] [𝑋 𝐹𝑒𝑂 𝑙𝑖𝑞𝑢𝑖𝑑 𝑋 𝑀𝑔𝑂 𝑜𝑝𝑥 ] ) 𝑐𝑜𝑟𝑒 -( [𝑋 𝑀𝑔𝑂 𝑙𝑖𝑞𝑢𝑖𝑑 𝑋 𝐹𝑒𝑂 𝑜𝑝𝑥 ] [𝑋 𝐹𝑒𝑂 𝑙𝑖𝑞𝑢𝑖𝑑 𝑋 𝑀𝑔𝑂 𝑜𝑝𝑥 ] ) 𝑟𝑖𝑚 (V. A. 3)
With 𝑋 𝑗 𝑖 being the mole fraction of component j in phase i. For the liquid components, the average compositions of the melt inclusions hosted in opx, studied in the blast (8 MI) and the Plinian phase (36 MI) (d'Augustin, 2021), with the ferric-ferrous ratio that was calculated using a calibrated equation from a study (Kress and Carmichael, 1991) and by knowing the temperature and fO2 (mgt/ilm pairs; Table V.2). Then, the variation in H2O content was deduced from equation ( 2), knowing the

∆ 𝐾 𝐷 𝐹𝑒 2+ -𝑀𝑔
between the cores and the rims of opx.

For the reverse-zoned last rims of the blast and Plinian phase, compared to their cores, the average variation of 𝐾 𝐷 𝐹𝑒 2+ -𝑀𝑔 was of 0.02 and 0.01, respectively, leading to a ∆𝐻 2 𝑂 𝑐𝑜𝑟𝑒-𝑟𝑖𝑚 of ~0.17 and 0.13 wt% H2O, showing a decrease in water content in the melt when the opx rims formed (∆𝐻 2 𝑂 𝑐𝑜𝑟𝑒-𝑟𝑖𝑚 = 𝑤𝑡% 𝐻 2 𝑂 𝑐𝑜𝑟𝑒 -𝑤𝑡% 𝐻 2 𝑂 𝑟𝑖𝑚 ).

For the normal-zoned rims of the blast and Plinian phase, compared to their cores, the average variation of 𝐾 𝐷 𝐹𝑒 2+ -𝑀𝑔 was of -0.03 and -0.05, respectively, leading to a ∆𝐻 2 𝑂 𝑐𝑜𝑟𝑒-𝑟𝑖𝑚 of ~ -0.31 and -0.46 wt% H2O, showing an increase in water content in the melt when the opx rims formed (∆𝐻 2 𝑂 𝑐𝑜𝑟𝑒-𝑟𝑖𝑚 = 𝑤𝑡% 𝐻 2 𝑂 𝑐𝑜𝑟𝑒 -𝑤𝑡% 𝐻 2 𝑂 𝑟𝑖𝑚 ).

Résumé

Ce chapitre qui aborde différents aspects de l'éruption de 2010-2013 du Kizimen est subdivisé en 4 sous-chapitres menant à l'estimation de durées de différents phénomènes pré-éruptifs. Les deux premiers portent sur les dynamiques pré-éruptives des magmas par l'analyse des cristaux naturels zonés (Chapters VI.1-2). Les deux derniers sous-chapitres portent sur l'étude des mélanges des magmas à travers deux expériences utilisant un appareil de mélange chaotique des magmas (Chaotic Magma Mixing Apparatus ; COMMA) et des échantillons naturels rubanés (Chapters VI.3-4).

Déchiffrer les échelles de temps des processus magmatiques dans le système d'alimentation magmatique menant aux éruptions volcaniques dans le passé peut fournir des informations pour la gestion des crises futures, surtout si une corrélation entre ces échelles de temps et les signaux de surveillance peut être établie. La dernière phase avant l'éruption est un épisode très court de « magma mingling » ou mélange inhomogène des magmas où toutes les poches de magmas sont mélangées suivies d'une remontée du magma en surface, produisant des laves rubanées. Les échelles de temps de cet épisode ont voulu être étudiées. Deux études expérimentales de mélange chaotique à haute température ont été réalisées grâce à une collaboration avec l'Université de Pérouse (Italie) avec deux objectifs :

-La première expérience visait à déterminer les cinétiques de mélange du premier événement de mélange de magma mafique avec le réservoir dacitique, produisant la partie andésitique en base de réservoir (pre-2009) -La seconde expérience de mélange visait à caractériser les échelles de temps de l'épisode de mélange des magmas juste avant l'éruption (compositions plus proches du système naturel) (Chapters VI.3-4). La première expérience montre la formation du magma hybride andésitique et l'homogénéisation du magma mafique en quelques heures. L'étude des échantillons naturels et la taille des cristaux montre que le mélange inhomogène des magmas s'est produit en quelques heures à quelques minutes avant l'éruption. La deuxième expérience, n'étant pas encore totalement exploitée pour la rédaction de ce manuscrit, permettra aussi d'estimer un temps sur les cinétiques de ce mélange.

CHAPTER VI.1: INSIGHTS INTO VOLCANIC UNREST BY CORRELATING PETROLOGICAL AND SEISMIC OBSERVATIONS AT KIZIMEN VOLCANO (KAMCHATKA, RUSSIA)

Insights into volcanic unrest by correlating petrological and seismic observations at Kizimen volcano (Kamchatka, Russia)

Lea Ostorero

Introduction

During volcanic unrest the information about the magmatic processes eventually leading to an eruption can be obtained from geochemical, geodetic, and seismic observations 1-4 . The latter are particularly valuable because of their depth and time resolution. An increasing number and intensity of volcanic earthquakes is considered the hallmark of an impending eruption 5 . While the correlation between the preparation of eruptions and the increased seismic activity is well established empirically, relating the latter to specific magmatic processes remains problematic. At the same time, the temporal evolution of magmatic processes in the plumbing system can be deciphered based on modern petrological and geochemical methods 6-9 . Therefore, a more direct link between seismicity and magmatic process can be obtained by comparing petrological to seismological records.

Within a reservoir, crystals grow from melt incorporating major and trace elements following thermodynamic and kinetic laws 10 . In silica-rich melt, orthopyroxenes (Mg,Fe)2Si2O6 are relevant candidates as their composition evolve from Mg-rich to Fe-rich composition, depending especially on melt temperature and composition, creating zoned crystals 6 . Diffusion in the zoned crystals is a timedependent process (Fick's law) acting as a clock for measuring timescales of a variety of magmatic processes 7,8,11,12 . Zoned orthopyroxenes (opx) in evolved mushy zones below volcanoes are thus time capsules allowing us to constrain the pre-eruptive timescales of magmatic processes 7,8,11,12 .

Timescales modelled using (Fe-Mg) interdiffusion chronometry applied to well-monitored eruptions, may be correlated to monitoring data (seismicity, changes in the gas chemistry and/or flux, deformation) and then have significant implications for volcanic eruption forecasting models 7,13-18 and a high potential for petrological monitoring of volcanoes 19 . Up to now only syn-eruptive investigations of magmatic processes in correlation with seismic activity have been performed for differentiated systems 17,20,21 compared to pre-eruptive correlations. To establish a temporal correlation between the magmatic processes and the seismic activity during a pre-eruptive re-activation of an arc plumbing system, we selected products from the first historical magmatic eruption of Kizimen in 2010-2013, in Kamchatka, involving differentiated magmas. This eruption is of particular interest because it occurred after several decades of quiescence and was well-monitored 22-25 . Seismic network operated by the Kamchatka Branch of the Geophysical Survey of the Russian Academy of Sciences resulted in an exceptional record of the pre-eruptive seismic re-activation. Furthermore, the petrological record in the opx crystals that is very sensitive to temperature fluctuations 20,26-29 has not been perturbed by previous volcanic activity. Therefore, this 2010Therefore, this -2013 Kizimen eruption provides us with a unique opportunity to study magmatic-seismic correlation and better understand the processes leading to intermediate to silicic eruptions. We report a good correlation between the petrological timescales and the pre-eruptive seismic monitoring 18 months prior the eruption. The petrological timescales are thus a valuable warning-clock for volcano monitoring, particularly in subduction settings.

Geological context and previous results about the 2010 eruption of Kizimen

The Kamchatka Volcanic Arc, in Eastern Russia, results from the subduction of the northeast corner of the Pacific plate at a rate of 8-9 cm/yr beneath the Okhotsk plate (which is connected to the North American plate) 30,31 (Figure 1a). This arc is the most productive arc on Earth (~310 millions of tons of magma emitted per year from the late Pleistocene to the Holocene) 32 , with more than 25 active volcanic centers. One of the world's largest clusters of subduction volcanoes, the Klyuchevskoy volcanic group, is located in the northern part of the arc, in the Central Kamchatka Depression 33,34 where the largest magmatic production is recorded 22,31,35 and where the volcanoes exhibit very sustained seismic activity (Figure 1a-b). The particularly strong volcanic activity in this region is likely caused by the hot mantle material brought by the around-slab-edge asthenospheric flow 36 and related crustal extension 37 . Nearly continuous occurrence of deep long-period earthquakes near the crust-mantle boundary 35,38 is likely related to the underplating of mafic magmas and widespread seismic tremors mark an active transcrustal magmatic system connecting different volcanoes of this group 39 . At the same time, welldeveloped magma chambers beneath some of most active volcanoes such as Bezymianny are well documented 40 .

Kizimen is located on the eastern margin of the Central Kamchatka Depression and its magmas might have similar origin to those from the Klyuchevskoy volcanic group. Because of the less dense coverage by seismic stations, the deep structure of the crust beneath Kizimen is not well known and the reported magmatic and seismic activity is linked to the upper crust. At the same time, the existence of crustal scale magmatic system in analogy with the Klyuchevskoy volcanic group cannot be excluded. Kizimen is a Late Pleistocene -Holocene volcano, with four cycles of eruptive activity identified (between 12,000-11,000 to 8,400 years ago, 8,400-6,400, 6,400-3,000, and 3,000 years ago to the present 34 ).

A possible phreatic eruption occurred in 1927-1928 34,41 but the first historical magmatic eruption of Kizimen occurred in 2010-2013, when magma volumes of 0.4 km 3 Dense Rock Equivalent (DRE) were emitted 41 (Figure 1b; Supplementary Figure 1). On 11 November 2010, explosive episodes generated pyroclastic density currents 22,41,42 , that were then followed by the extrusion of a lava dome which gradually formed a ~200-m thick lava flow-41 (Figure 1b; Supplementary Figure 1). Numerous block and ash pyroclastic density currents were produced as a result of the gravity collapse of the lava dome and flow front. Dacites, silica-rich andesites and banded lavas, formed by an incomplete magma mixing event between two magmas were emitted 41 (Figure 1b-f; Supplementary Figure 1 and Supplementary Data 1). For this eruption, a model of the pre-eruptive dynamics proposed that the emitted dacite came from the differentiation of a basaltic magma 41 . This basaltic magma was not erupted at the surface 41 . Several models have assessed the presence of a stable silicic reservoir under Kizimen, periodically refilled by magmatic injections 41-43 . This silicic magma chamber is located between 5-11 km depth (based on earthquakes depths and amphibole and opx compositions 41 , which is in agreement with a study on older eruptions from Kizimen with storage pressures of magmas between 1 to 3 kbars (up to 10 km depth) 42,43 ). Another seismic zone from the ground surface to ~5 km has been recognized

during the 2010-2013 eruption and probably shows the conduit in which magma ascended to the surface 41 . An intensive seismic swarm 22,41 was recorded in 1963, with earthquakes depths reaching down to more than 35 km but without leading to an eruption 41 . This seismic swarm is interpreted as a mafic injection from depth, as the samples from the 2010 eruption that followed showed resorbed olivines and as it occurs in a similar geologic and tectonic situation at the volcano Tolbachik in the Klyuchevskoy group, where intrusion of dikes of mafic magmas generate multiple earthquakes at depths of up to 30 km 41 . An episodic increase of the fumarolic activity also followed the occurrence of large shallow earthquakes of this seismic swarm (magnitudes of 6.2 and 5.8) 34 . However, due to the high viscosity of magma in the reservoir, a long time period can exist between the input of basic magma and the following eruption 41 . This mafic injection led to magma hybridization of the lower part of the dacitic reservoir, forming an andesitic magma at the base of the magmatic reservoir 41 . A large portion of the dacitic magma was not affected by this mafic recharge, leading to the formation of a zoned reservoir.

A seismic crisis started ~1.5 years before the eruption with earthquakes located between 11 km and the surface 22,23,41 . In addition, deformation of the volcanic edifice was detected by Interferometric synthetic aperture radar (InSAR) 24 . The seismic data and deformation by InSAR were interpreted as the injection of a dike of 15 km long and about 20 cm thick in the magmatic reservoir 24 . As no seismicity below 12 km has been recorded, this interpretation is thought not to fit the existing geological and seismological data 41 . Prior to the eruption of 2010, another mixing event occurred between the dacitic and andesitic part of the reservoir due to some degassing of the dacitic magma leading to magma convection, with incomplete mixing between the two magmas 41 . Kizimen is crossed by an active tectonic fault, through which volatiles are escaping to the surface. This leakage is visible as a persistent fumarole on the NW flank of the volcano 41 , this long-term passive degassing could be responsible for the depressurization of the magma chamber and increase the viscosity of the magma 44,45 (Figure 1b). The temperatures of the fumaroles at the summit of Kizimen have been measured from 15-31 August 2009 and a temperature of 340 °C was reached close to the central mouth of the outermost fumaroles on Kizimen which is higher than measurements prior to the onset of the unrest (maximum of ~270 °C) 46 . However, this temperature increase of the fumaroles is not thought to be a consequence of seismic activation 46 . One month before the eruption, visual signs of reactivation were observed with thermal areas and two new large fumaroles that appeared (16 October 2010) 22,41 .

Results

Magmas are differentiated, with whole rock compositions in the andesitic and dacitic field (Figure 1c )). Three compositional groups are identified: the first one corresponds to all the cores of the opx (both unzoned and zoned opx of the dacite and andesite; En62-64), then a group for the normal-zoned rims of the opx of the dacite (En44-54) and finally a group for the reverse-zoned rims of the opx in the andesite (En64-73) (Figure 2; Supplementary Figure 45and Supplementary Data 2). Harker diagrams also highlight that the unzoned opx and cores of the zoned opx have the same compositions (Supplementary Figure 5).

Timescales

Of 112 profiles on zoned crystals, 26 were suitable for diffusion modelling for the dacite, as well as 62 for the andesite (Supplementary Figure 678; Supplementary Note and Data 2). The temperatures Supplementary Figure 6; Supplementary Data 2). For most of the short timescales (< 2 years), the positive error associated to the timescale is in the order of 1 year and the negative error is around 100 days (Supplementary Figure 6; Supplementary Data 4). The young timescales have lower uncertainties than the timescales superior to 1 year before the eruption (Figure 3e).

The temperatures used to model the timescales were determined using the compositions of magnetite/ilmenite pairs which can show the variability of the temperatures in the reservoir. Varying the temperature and using the one determined in opx in the andesitic part of mingled samples by another study 41 implies shorter timescales for the andesite (divided by 2, with T = 875°C), showing the importance of the choice of the temperature in the timescale modelling. As no temperature was determined in opx of the dacite in the previous study 41 and no melt inclusion with a sufficient size was found in our samples, the temperatures determined on the magnetite/ilmenite pairs were used.

On Figure 3, considering only the major En changes of the zoned opx (normal zoning for the dacite (from En62-64 to En44-54) and reverse zoning for the andesite (from En62-64 to En64-73); Supplementary Data 4), all the zoned opx give timescales equal to or below 1.5 years. The cumulative number of timescales recorded mostly increase 1 year before the eruption (Figure 3d). So, 1.5 years prior to the eruption of 2010, after ~50 years of dormancy, the opx in the dacitic and andesitic magmas begin to record a progressive magmatic disturbance accompanied by changes in the storage conditions, with different processes as their most external zonations are not the same (Figure 3b andd; Supplementary Figure 234). One year before the eruption, these magmatic processes become more dynamic (as the curve of the cumulative timescales is steeper than before; Figure 3d). In the andesite, reverse zonations (~80%) in the zoned opx reflect heating processes 26,27 or degassing 53 and normal zonations (~20%) reflect cooling processes 15,17,26,27 or a high water content in the melt 53 (Figure 2; Supplementary Figure 234). In the dacite, dominant normal zonations (93%) principally register cooling processes 15,17,26 or a high water content in the melt 53 (Figure 2; Supplementary Figure 234).

To explain the reversed-zoned rims of the andesite, four hypotheses can be evoked: i) the injection of a mafic magma at the base of the reservoir, 1.5 years prior to the eruption.

This mafic magma would have been aphyric as no opx have a different En core composition. But as no deep earthquakes (> 11 km) have preceded and accompanied the 2010 eruption, this event is not likely to have been triggered by input of basic magma from greater depth 41 , unless magma ascent occurred through previous pathways, explaining the lack of deep seismicity. But if the last mafic injection occurred in 1963, it is likely that fracturing of the rocks needs to happen for the ascending magma, after ~50 years, generating seismicity. No enclave of basaltic or basaltic andesite compositions were found, unlike many of the earlier eruption products where mafic recharge before the previous eruptions have been inferred 42,43 . Furthermore, the melt inclusions data are in equilibrium with their hosting opx (Supplementary Figure 9) and they are in the same domain as the residual glasses (Supplementary Figure 10). The melt inclusions do not show different

CaO or Al2O3 values compared to the En content compositions (Supplementary Figure 10). So, either the mafic magma only heated the andesite without mixing with it (which is unlikely, with the chaotic nature of mixing processes 54-56 ), generating 30% of zoned opx or it mixed in a part of the reservoir without reaching the surface during the eruption.

ii) A convective event in the andesitic part of the reservoir, zoned in temperature, between the walls and cores of the reservoir, generating the reverse-zoned rims in the andesite.

However, this convective magma mixing process is unlikely to occur without affecting the dacite, as only one reservoir, or a reservoir with linked magma pockets, is more likely to be present at Kizimen (depths of 5-11 km 41 ; locations of earthquakes and amphibole and orthopyroxenes compositions), even if magma pockets can remain isolated.

iii) A partial mixing in the reservoir at the contact between the andesitic and dacitic magmas, with crystal exchanges. Some opx from the dacite could have migrated in the hotter andesite, generating reverse zonations in the opx (rims of En62-64) (Figure 4b). In parallel, some opx from the andesite could have been in contact with the dacitic magma generating normal zonations (rims of En55-61) (Figure 4b). This mixing process only affected part of the two magmas in contact, explaining that 70% of the opx in the andesite and dacite are unzoned.

iv)

Degassing processes in the andesite, with a decrease of the water content in the melt, which favors Mg-rich opx (producing reverse zoning) 53 . Indeed, series of H2O fluid-saturated phase-equilibrium experiments showed that the Fe 2+ -Mg exchange coefficient KD, between a rhyolitic melt and minerals, such as opx, increases as a function of dissolved water in the melt 53 . An increase of the Mg# of the opx was observed when the H2O content in the melt decreases, which means that Mg-rich opx (reverse zoning) crystallizes at low water contents whereas Fe-rich opx (normal zoning) are favored at high water contents 18,53 .

Normal-zoned plagioclase (sodic rims) generally accompany reversely zoned opx in the case of loss of H2O during degassing 53 , of the andesite was estimated at ~0.06 and was used to estimate the variation in H2O content in the melt, from the core of the opx to the rim. Using an experimental relationship 18,53 , this ∆ 𝐾 𝐷 𝐹𝑒 2+ -𝑀𝑔 corresponds to a ∆𝐻 2 𝑂 𝑐𝑜𝑟𝑒-𝑟𝑖𝑚 of ~ 0.7 wt% H2O, showing a decrease in water content in the melt when the opx rims formed (∆𝐻 2 𝑂 𝑐𝑜𝑟𝑒-𝑟𝑖𝑚 = 𝑤𝑡% 𝐻 2 𝑂 𝑐𝑜𝑟𝑒 -𝑤𝑡% 𝐻 2 𝑂 𝑟𝑖𝑚 ) (see Supplementary Note). However, the initial H2O content of the melt was not estimated, as the few melt inclusions studied here were too small to be analysed with the Secondary-ion mass spectrometry (SIMS), so the initial dissolved water composition is not known and cannot be compared with the experiments, that studied the Fe2+-Mg KD variations as a function of dissolved water in the melt (from 2.7 to 5.6 wt%) 53 . Still, we calculated an estimate of the H2O content (2.4 wt% H2O), using only one melt inclusion in a plagioclase from the dacite (plag/liq thermometer 57 ; and hygrometer 58 (Supplementary Note and Supplementary Data 1).

Provided that all these estimates are correct and based on only one melt inclusion, the H2O content decreased from 2.4 wt% to 1.7 wt%, which is outside of the water contents investigated in the experimental study 53 . We are then not able to conclude on this hypothesis, due to the scarcity of our melt inclusions data.

The hypothesis of magma mixing between the dacitic and andesitic magmas is then more likely, given the points discussed above. But if it explains the normal and reverse zonations in the opx of the andesite, it does not explain the normal zonations in the opx of the dacite. The high viscosity of the dacitic reservoir probably prevented the magma mixing of all the dacite with the andesitic part of the reservoir 41,59 .

The normal zonations in the opx of the dacite can be explained by cooling 15,17,26,27 or a high water content in the melt 53 . As only 30% of the opx in the dacite are zoned and that normal zoning is dominant (93% of the zoned opx; Supplementary Figure 2), an increase in the H2O content in the melt by differentiation could have led to the development of normal zoning in the opx 53 . Some movements of some opx of the dacite in colder areas in the dacitic part (for example with the walls of the dacitic reservoir, which are colder than the core), could also have led to the formation of the normal zonations (Figure 4b). ∆ 𝐾 𝐷 𝐹𝑒 2+ -𝑀𝑔 between the cores and normal-zoned rims of the opx in the dacite is equal to ~ -0.05, corresponding to a mean water content change of ~ -0.5 wt% H2O (see Supplementary Note).

This means that there was a water content increase in the melt from the core to the rims of the opx (from 2.4 wt% to ~3 wt% H2O), which is in agreement with the experimental study, where normal zoning is favored at higher water content in the melt 53 . However, in the light bands of the banded samples, normalzoned plagioclase are dominant 41 , which is not coherent with the presence of normal-zoned opx in the dacite, according to the experimental study 53 . Further investigation of the plagioclase in the dacite is then also needed. The hypothesis of some partial internal mixing within the dacitic pocket with colder areas is then preferred to explain the normal zonations in the dacite (Figure 4b).

A period of unrest then follows, while the magmas reach an eruptible state 41 , with magma mixing processes still occurring, up to the time of the eruption.

Finally, magma mingling occurred between the andesitic and the lower dacitic part, during ascent shortly prior to the eruption, forming the banded andesites and dacites (Figure 4c). Magma mingling events between different magma types are also recognized as very short events prior to an eruption, as shown combining magma mixing experiments, numerical simulations and statistical analysis before eruptions in the Campi Flegrei volcanic system in Italy, where mixing processes can be of the order of a few days to a few minutes 56,60-62 .

The eruption trigger could then correspond to the magma mixing process beginning 1.5 years before the eruption and increasing progressively until it reaches an eruptible state, with the last magma mingling phase and then the eruption. Another possibility for the eruption trigger could be a decompression of the dacite, generating the last magma mingling phase leading to the eruption.

However, as no explosive products from the first phases of the eruption were found, it is difficult to conclude on a scenario for the eruption trigger.

CO2 flushing could also participate in triggering the eruption, i.e. interaction between carbonic fluids and H2O-rich magmas stored in the upper crust 63 . Due to fluid saturation and magma ascent, release of carbonic fluids can occur, increasing the magma volumes and generating overpressure that can destabilize the magmatic system, leading to the eruption 63 . A way to test this hypothesis is to study melt inclusions H2O and CO2 contents, however, as said before, no H2O and CO2 content measures were carried out in the samples, as the melt inclusions were too small (only one melt inclusion in a plagioclase gave a water content of 2.4 wt% H2O; plag/liq thermometer 57 ; and hygrometer 58 ) (Supplementary Note and Supplementary Data 1). Data from the literature of the same eruption in the banded lavas give H2O content between 5-5.7 wt% H2O in the dacite and 4.9-5.6 wt% H2O in the andesite (amphibole thermometers 41,64,65 ) but no CO2 content have been measured. From an older eruption, in 7550 14 C BP, H2O content are of 0.5-3.5 wt% in melt inclusions in plag and mgt, 0.6-4.5 wt% H2O in amphiboles and 1.6-5 wt% in opx and CO2 of 62-9 260 ppm (with 2 over 46 above 4000 ppm) 66 . As data from another eruption is not relevant to study the hypothesis of CO2 flushing, we are not able to conclude on this hypothesis.

The distribution of timescales obtained on the opx crystals can be correlated with the seismic activity (Figure 3). For this goal, we used the catalog of volcanic earthquakes produced by the Kamchatkan Branch of Russian Geophysical Service. The completeness magnitude for this catalog is difficult to estimate and can be variable in time depending on the number of stations and meteorological conditions. Therefore, instead of reporting the number of recorded earthquakes (that is strongly sensitive to this completeness magnitude), we use the cumulative seismic moment that is dominated by strongest earthquakes (Figure 3a-d). Also, this quantity is closely related to the overall energy seismic release and therefore has a physical relationship with ongoing processes within a volcano. Because

Kizimen remained quiescent for decades and is located far from other volcanoes, this catalog is exclusively related to the 2010 eruption (and to the magma recharge potentially in 1963 that did not lead to the eruption 41,67 ). Comparison of the timescales modelled with the cumulative number of earthquakes is also shown (Supplementary Figure 11), with an increase of the number of timescales that record magma mixing processes closer to the time of the eruption, even if the number of earthquakes is increasing faster than the rate of crystal clocks starting. This could be due to the increase of seismic unrest near the time of the eruption, with a relatively lower rate of magma mixing. An inflection of the cumulative seismic moment and number of timescales can be evidenced, from ~January 2010-November 2010 (~one year before the eruption; Figure 3d; Supplementary Figure 11). This inflection can correspond to an intensification of the magma mixing processes (between the dacite and the andesite). As this process becomes more dynamic, the opx rims are formed more and more rapidly.

80% of the modelled timescales cover the period between April 2009 and the onset of the eruption on 11 November 2010 and 60% from January 2010 to November 2010 (Figure 3b andd). This increase in number of petrological timescales coincides with the onset of the seismic crisis recorded ~1.5 years before the eruption 22,41 , with earthquakes between ~3 to 13 km from April 2009 to January 2010 and then becoming more numerous and shallower from January 2010 to the eruption (Figure 3a andc, Supplementary Figure 8). This increase of the timescales is linked to magmatic reactivation, as discussed before, with magma mixing at the contact between the two magmas. Magma mixing probably generated an overpressure in the reservoir at the origin of the seismicity recorded during the same time period (Figure 3a-d), April 2009-November 2010. As the timescales are also more numerous 1 year before the eruption, coinciding with the increase of the seismic moment around that same time (January 2010; Figure 3c-d), increase in the magma mixing dynamics between the andesite and dacite could induce bubble nucleation (and overpressure), leading to an increase in the seismicity 68 . The cumulative number of timescales and cumulative seismic moment are perfectly overlaid around April 2010 (8 months before the eruption; Figure 3c-d), probably because of the increase in the magmatic processes near the onset of the eruption. Finally, the very short seismic and timescales signals close to the eruption (in the order of hours to a few days; steeper curve in Figure 3c-d) can correspond to the magma mingling event, prior to the eruption on 11 November 2010, as evidenced before several explosive eruptions 69 .

We note that the apparent migration of the seismicity around January 2010 becoming shallower (Figure 3a andc), between -5 km to +3 km is likely due to the modification of the software used for the analysis of the earthquakes 22 . While this modification might affect the number of detected small earthquakes its influence on the cumulative seismic moment is negligible (Figure 3b andd). Progressive ground deformation was also detected before the eruption, from September 2008 to September 2010 (InSAR data) 24 . These types of signals could be partly explained by the increase of fluid pore pressure along the plane of the tectonic fault 41 and be related to magma mixing and movements 13,14 . The increase of temperatures of the fumaroles measured in August 2009 46 , are compatible with the magmatic disruptive processes in the reservoir (Figure 4).

The opx chronometer is thus successfully used here to determine the links between pre-eruptive magma dynamics in the plumbing system and pre-eruptive monitoring data at the surface for an eruption involving differentiated magmas. This chronometer shows the reactivation of the magmatic feeding system of Kizimen 1.5 years before the eruption, well correlated with the monitoring signals recorded on the surface from April 2009 to the time of eruption. Most of the studies investigating petrology and monitoring signals are carried out in mafic systems, using olivines, such as Mt Etna (Italy) 13,15 , Kilauea (Hawaiʻi) 70 , Eyjafjallajökull (Iceland) 14 , Piton de la Fournaise (La Réunion) 71,72 , El Hierro (Canary Islands) 2,73 , and arc volcanoes such as Shishaldin volcano (Alaska) 16 , Merapi (Indonesia) 74 (using clinopyroxenes), Irazú (Costa Rica) 75 or Llaima (Chile) 76 . These studies obtained mixing-to-eruption or recharge events of the order of a few days to months (3-11 months 2,14-16,71,73,76,77 ) to years before the eruption (1-2 yrs) 13,70,74,75 . Compared to monitoring signals, it depends on the eruption, but overall, the petrological timescales match well with seismic precursory activity months to days before the eruption 2,16,71,73,76 or with the GPS displacements 14 . The timescales obtained on arc volcanoes such as Merapi or Irazú agree with the ones found on Kizimen (2 and 1.1 yrs in average, respectively), which are also consistent to monitoring signals (seismicity, deformation) 74,75 .

Only three studies have used opx chronometer, but to investigate syn-eruptive monitoring 17,20,21 . On Mount St-Helens correlations were made with some eruptive phases during the 1980 eruption but no correlation was made with seismicity before the beginning of the eruption 17 . According to the diffusion timescales modelled, timescales of the order of 1 year before the eruption in 1980 were found 17 , without being correlated to the seismicity that began 2-3 months before the eruption. On Ruapehu (New Zealand) 20 as on Bezymianny (Kamchatka) 21 , the quasi-permanent high seismic activity of these volcanoes makes it difficult to clearly correlate timescales and seismicity, but magma perturbations occurred months to years before eruption. One recent study has used opx to decipher the processes and pre-eruptive conditions that led to the 1990 Kelud basaltic andesite eruption (Indonesia) and showed a relationship between the timescales (pre-eruptive fluid accumulation a few weeks to months prior to the eruption) and the change of hydroacoustics and water lake temperature but the relationship with seismic swarms was more diffuse 18 .

A comparison between the diffusion timescales found by the previous studies coupling diffusion modelling with monitoring signals with Kizimen ones modelled here is shown in These eruptions are thought to have been preceded by magma mixing or magmatic injections before the eruption, which are the scenarios investigated here at Kizimen. These correlations could have implications in terms of monitoring, with unrest periods of more than a year that could lead to eruptions.

This long period of unrest has been observed for high silica systems that have experienced long quiescent time where an eruption can occur after a long alert period (from days to years) 78 .

More broadly, a contrast between pre-eruptive timescales obtained on basaltic and silicic volcanoes can then be highlighted, with short timescales obtained in mafic systems compared to more evolved systems 7 . Indeed, magmatic perturbations prior to eruptions emitting basalt or basanite are generally occurring from days to months before the eruption (olivine studies) 13-16,49,50,70-73,79-85 . For more evolved magmas (basaltic andesite/andesitic/dacitic magma; on olivines, opx, cpx, plag or apatites) 18,20,21,28,86 , timescales are in the order of a few years 21,74,75 to decades or more 26,87-91 , even if some eruptions have shorter pre-eruptive timescales (~1 year). For dacitic to rhyolitic magmas, using the same crystals as for the previous systems but also quartz, timescales are between one year to longer timescales, such as centuries 17,26,29,48,90,92-99 and also correspond to magma recharge or heating experienced by the magmas.

The timescales found for Kizimen 2010 eruption are then in a rather young range compared to these systems.

To conclude, acquiring this type of petrological data on recent and past eruptions on a given volcanic system is of great importance. Petrological disturbances in a storage system are synchronous with geophysical or geochemical signals followed by monitoring institutions. The strong temporal correspondence between timescales and monitoring data (seismicity, deformation...) that we found at Kizimen (magma mixing that occurred in the reservoir 1.5 years before the eruption, correlated to the onset of a seismic crisis), shows the possible application of this approach on volcanoes whose past eruptions have not been monitored but on which it is possible to estimate timescales by petrology. It is often difficult, particularly on intermediate to silicic volcanic systems, to determine when the first warning signs are recorded and how long it will take before a possible eruption. Moreover, if one can establish a systematic pattern in the timescales for several past eruptions, as it has been established at Montagne Pelée, Martinique 100 , it can be applied more easily for a future eruption. These results indicate that this type of data can be used to estimate the timescales between the first precursory signals and the beginning of an eruption in case of reactivation of a volcano.

Materials and Methods

Sampling on the field: andesites, dacites and banded samples

Samples were collected in the pyroclastic deposits of the 2010-2013 eruption (Figure 1a) during a three-week field mission which took place in August 2019. Samples coordinates can be found in Supplementary Table 1.

Sample preparation

The unaltered dacitic and andesitic samples were crushed (up to 3 mm). One half was then crushed again into fine powder for whole rock analysis while the other half was sieved into different size fractions (710-1000, 500-710, 355-500, 250-355 and 125-250 μm fractions)). These fractions were washed in ultrasonic bath and dried at 80 °C for 48h. They were then observed under the binocular microscope to select the fractions in which the crystals were the more abundant and automorphic. Opx, plagioclase (plag), magnetites (mgt) and amphiboles (amph) were then handpicked under the binocular microscope in the different size fractions. They were mounted in epoxy resin and polished up to 0.3 µm to the middle part of the crystals. Opx were oriented with c-axis in a north-south direction as they have been used for intracrystalline diffusion modelling 20,26,28,90,101 . Before scanning electron microscopy (SEM) or electron microprobe micro-analysis (EPMA) investigations, selected mounts were all carboncoated 90 .

Textural observations: Scanning Electron Microscope

Opx were observed under Scanning Electron Microscopes (SEM): Zeiss Supra 55VP (Sorbonne Université, Institut des Sciences de la Terre de Paris, ISTeP, Paris) and the Carl Zeiss EVO MA10 SEM at the PARI platform at the Institut de physique du globe de Paris (Université Paris Cité) using an acceleration voltage of 20kV and a beam current of 8 nA. An identification of possible chemical zonations in the crystals (grey levels of different intensities) was possible in order to position the rim-core profiles to be carried out with the electron microprobe for opx. Proportions of zoned/unzoned opx crystals were determined on all the crystals mounted by looking closely at the SEM images.

High resolution backscattered electron (BSE) images (mag x755) were taken for diffusion modelling and intercalibration with chemical profiles of opx with the same parameters mentioned above, a high integration line (N=7, 7 integrations of each image line to reduce the impact of noise) and a dwell time per pixel of 48 µs.

Melt inclusions in crystals and residual glass areas were also identified using SEM images. Melt inclusions that were of sufficient size for electron microprobe analysis (> 6 µm), round in shape, suggesting the absence of crystallization after trapping in the crystals and not containing bubbles (which could indicate outgassing) were selected. The selected glasses areas had a size greater than 10 µm and without microcrystallization.

Compositional analysis

Whole rocks

Whole rock major-element compositions were analyzed by ICP-OES at Centre de Recherches

Pétrographiques et Géochimiques (CRPG) in Nancy 102 .

Point measurements

Opx crystals were analysed for major elements (Si, Al, Ca, Mg, Na, K, Ti, Fe, Mn and P) by electron microprobe micro-analyzers with an acceleration voltage of 15 kV, a beam current of 10 nA and a focused beam of 2 µm (CAMECA SX-Five and SX-100; Service Camparis, Paris). Counting time on peak and background was set at 10 s for all elements, except for Fe and Mg with 80 s. The core to rim compositional profiles in zoned opx had a 2 µm step and an average length of 100 µm (~4 hours per profile). Four points were measured across the unzoned crystals. They were acquired perpendicular to the long axes of the opx and away from the corners to avoid three-dimensional effects such as growth 29,90,101,103 .

For compositional analyses in magnetite/ilmenite pairs to estimate the temperature, the counting time was set at 30 s for Fe, Ti and Al and 10 s for the other elements.

For the measurement of major and minor elements in the residual glasses and melt inclusions of the samples, an acceleration voltage of 15 kV was used, with counting times varying from 10 to 25 s depending on the element and a beam size of 6-10 µm to avoid alkaline migration 104-106 .

Analytical errors on major and minor elements were calculated to be of 0.7 wt% for Si, ~0.1 wt% for Na, Al, Ca, K and Ti, 0.07 wt% for Ni, 0.29 wt% for Mg, 0.39 wt% for Fe, 0.14 wt% for Mn, 0.04 wt% for Cr and P and 0.02 wt% for F and Cl (for melt inclusions and residual glasses).

Timescales of magmatic processes

To investigate the changes in crystallization conditions in time and date the magmatic processes, interdiffusion timescales were modelled following the shapes of the concentration's profiles in opx and width of the diffusion zone between the plateaus. Using second Fick's law, if the initial conditions, boundary conditions and diffusion coefficient of the elements are known, fitting of the profiles can be used to obtain the timescales 6,7,13,15,49 . The diffusion coefficient depends on parameters such as chemical composition (Xi; molar fraction of the mineral constituent element), temperature (T in Kelvin), pressure (P in Pa), oxygen fugacity (fO2) and water fugacity (fH2O) 6,7 . Studies have modelled the timescales on different types of crystals (olivines, plagioclase, clinopyroxenes, orthopyroxenes, quartz…) using different diffusion coefficients and modelling 7 . Some programs exist to model timescales based on EPMA data, for example in olivines (Diffusion Process Analysis; DIPRA 107 ), or on BSE images in clinopyroxenes (on greyscale data using the Non-Isothermal Diffusion Incremental Step model;

NIDIS 108 ) or in orthopyroxenes 28 .

Modelling of the timescales associated to the zonations in the zoned crystals was done by using the method developed in other studies 85,101 , using the parametrization of the interdiffusion coefficient D of Fe and Mg elements in opx 109 . The Fe-Mg interdiffusion profiles were modelled in one dimension, across the c-axis, parallel to the b-axis of the crystals 85,101 . D depends on the composition of the opx, temperature, pressure, oxygen fugacity and water fugacity 6,26 . The interdiffusion coefficient we used here is formulated for a temperature between 500 and 800°C, an oxygen fugacity of IW + 0.8 log units above the IW buffer (IW: Iron-Wüstite buffer) and XFe ~0.10-0.50 109 (XFe is the molar fraction of the ferrosilite component). The parametrization of D in our model was done with a range of compositions of opx that includes those of this study (En44-73; XFe = 0.23-0.53; Supplementary Figure 4) 109 , the temperature of the magmas are slightly above the range used in the parametrization of D, which is extrapolated at our temperatures. In the experimental study of D we used to model our timescales, an fO2 dependence similar to olivine with an exponent of 1/6 was hypothesized 109 . This fO2 dependence has been incorporated in the formulation of D in other works 96,101 but this dependence has never been verified experimentally for opx with En contents of our dataset 110,111 . For En-rich opx of the mantle (En98 and En91), between 870-1100°C and oxygen fugacity controlled (fO2 =10 -11 -10 -7 Pa), an fO2

dependence of an exponent between 0 and 1/20 was found to be appropriate from their data 111 . The estimate for DFe-Mg used here 109 , when extrapolated to higher temperatures of the study on En-rich opx without considering any fO2 dependence appear to be consistent with their dataset 111 . The oxygen fugacity dependence on Fe-Mg interdiffusion coefficient is still under discussion as incorporating it can give longer timescales 48 but no study has yet examined the appropriate En compositions of the opx and appropriate oxygen fugacity ranges. Therefore, the parametrization of the interdiffusion coefficient D used is the one determined on opx compositions close from our dataset without an fO2 correction 109 , as done in previous studies 26,90 . D is defined as follows in equation ( 1) 109 :

logD = -5.54 + 2.6X 𝐹𝑒 - 12530 𝑇 (1)
The main assumption of the model is that the initial profile between two zones follows a step function, which is modified by diffusion to form sigmoidal concentration gradients 112 (Supplementary Figure 8). This steep profile becomes broader with time, due to interdiffusion of the different elements. With the eruption, the diffusion is stopped. Maximum timescales are then estimated 112 .

Profiles with smooth compositional variations of sigmoid shape were studied and those displaying growth dominant profiles were not considered (Supplementary Note; Supplementary Figure 7; Supplementary Data 2). These profiles display linear trends or curves and bumps on the profiles and are not consistent with a diffusion sigmoid calculated for the given initial compositional contrast 26

(Supplementary Figure 7). The Al and Ca gradients were compared to the Mg#, as Al and Ca diffuse slower than Fe-Mg 48,101,113,114 (Supplementary Figure 7; Supplementary Data 2).

Diffusion timescales were modelled using a spreadsheet done by D. J. Morgan (now published as a demo version for olivines 85 .

Examples of modelled profiles can be found in Supplementary Figure 8 and best fit profiles are shown for all modelled profiles in Supplementary Data 2.

Before using the diffusion model, an intercalibration was done between profiles of greyscale values determined from BSE images of zoned opx using the ImageJ software (https://imagej.nih.gov/ij/; version 1.52a) and with EPMA profiles done at the microprobe in the same crystals. We followed the methodology developed in two other studies 85,101 .

The grayscale profiles were done on high-resolution north-south oriented images taken with SEM (spatial resolution of 755 nm), and were averaged over a certain width to reduce the noise (between 20

to 50 µm, depending on the crystal rim size) 26,90 . The Mg number is the parameter that seems to best control the grayscale of the images 52,85 . Best fit diffusion profiles to the calibrated BSE/EPMA profiles were then modelled using the model spreadsheet 26,52,85,101 to obtain the timescales of interdiffusion. As the diffusion coefficient used is composition-dependent, the shape of the curve of the profile was determined by the degree of compositional contrast but the curve width is dependent on time 26,85 . A database of simulated diffusion profiles obeying composition-dependent diffusion under a 1D diffusion geometry was generated using a finite-difference method 26,52,85,101 . As all 1D diffusion profiles are selfsimilar in time for any specific compositional dependence, the curves are identical in their proportional curvature for initial boundary conditions (such as T, fO2, anisotropy and compositional contrasts) 85 .

Thus, if the boundary conditions and the initial condition are known for a specific measured diffusion profile, the correct diffusion profile shape can be determined from these simulated diffusion curves as the one with the same degree of compositional contrast and asymmetry characteristics 85 . When the appropriate profile has been identified in the database, a linear stretch is applied along the distance x axis to fit the data 85 . The spreadsheet measures the distance between the 20 th and 80 th percentile in the measured profile and then scales the appropriate model curve by the correct n factor before overlaying them (centred on the 50 th percentile of both curves as a reference point in the centre of the diffusion range) 85,101 .

The timescale corresponding to the zonation is then obtained, as the scaling factor is an expression of crystal composition, the diffusivity used for the library of the profiles, the diffusivity in the crystal and the ratio between the timescales of the database and the timescale associated with the zoning in the crystal 26 (with the timescale of the crystal being solved as the only unknown) 85,101 .

It is possible to write that 𝑡 𝑠𝑎𝑚𝑝𝑙𝑒 is:

𝑡 𝑠𝑎𝑚𝑝𝑙𝑒 = 𝑛 2 𝐷 𝑚𝑜𝑑𝑒𝑙 𝑡 𝑚𝑜𝑑𝑒𝑙 𝐷 𝑠𝑎𝑚𝑝𝑙𝑒 ( 2
)
Where n is the stretch factor 85,101 . The goodness of the fits is done by fitting the profiles by eye 26,90 .

The determination of the temperature, the diffusion coefficient and its measurement are the main sources of uncertainties 6,112 . Uncertainties on the calibration of measurements and point spacing during EPMA analyses, as well as the resolution of SEM images are also sources of uncertainties (pixel size, uncertainty on grayscale values defining a plateau) 90 . Overall, the measurement of the diffusion coefficient and temperature represent the main sources of uncertainty and have a logarithmic effect on final calculated timescales 6,26,52 (Figure 2; Supplementary Figure 6). Here, a Monte Carlo simulation was used to estimate the integrated time uncertainty from the different sources of uncertainty (on T, D0, Ea and resolution of grayscales values of SEM images), giving asymmetric uncertainties, with a larger error bar on the positive error and a smaller error bar on the negative error 26 .

Estimation of earthquake magnitudes

The Kamchatka Branch of Russian Geophysical Survey uses the following magnitude 𝑀 scale in equation ( 3):

𝑀 = 𝑙𝑜𝑔 10 𝑈 + 𝑓(𝑡 𝑆-𝑃 ) (3)
With 𝑓(𝑡 𝑆-𝑃 ) expressed in equation ( 4):

𝑓(𝑡 𝑆-𝑃 ) = { 1.42 𝑙𝑜𝑔 10 (𝑡 𝑆-𝑃 ) + 0.72, 𝑡 𝑆-𝑃 < 7𝑠 2.44 𝑙𝑜𝑔 10 (𝑡 𝑆-𝑃 ) -0.14, 𝑡 𝑆-𝑃 ≥ 7𝑠 ( 4
)
where U is the maximum S-wave ground velocity in micron/sec and 𝑡 𝑆-𝑃 is the arrival time difference between S and P waves. The seismic moments 𝑀 0 (in N m) are approximately estimated from magnitudes based on the moment magnitude equation in equation ( 5):

𝑀 0 = 10 (1.5 𝑀 + 9.05) (5)

Data and materials availability

The seismic catalog used for the analysis is available from the website of the Kamchatka Branch of the Geophysical Survey of Russian Academy of Sciences (http://sdis.emsd.ru/info/earthquakes/catalogue.php).

Figures

Figure 1. cm/yr 30,115,116 all data was available) and error bars correspond to the minimum and maximum timescale modelled 13-15,17,20,71,72,75,76 . For some studies, not all the timescales modelled were provided in the articles (or only the timescales ranges were provided) 2,21,70,73,74 so the values taken were the minimum and maximum timescales given in Figures or in the text of these studies. The comparison between these studies is summarized in Supplementary Data 5.

plag/liquid thermometer (equation 24a) 57 , from an older eruption of this volcano, in 7550 14 C BP: 0.6-4.5 wt% H2O and CO2 62-249 ppm 66 .

∆ 𝐾 𝐷 𝐹𝑒 2+ -𝑀𝑔

estimations. An increase of the Mg# of the opx was found when the content of H2O in the melt decreases, for a series of H2O fluid-saturated phase-equilibrium experiments 53 . An experimentally calibrated relationship between the 𝐾 𝐷 𝐹𝑒 2+ -𝑀𝑔 and the H2O content was found by the same authors 53 , comparing their data with another experimental study 121 . Based on the available data, we found this equation doing a linear regression:

𝑲 𝑫 𝑭𝒆 𝟐+ -𝑴𝒈 𝑜𝑝𝑥 -𝑙𝑖𝑞𝑢𝑖𝑑 = 0.0979 (𝑤𝑡% 𝐻 2 O) -0.0007 (2)
We then calculated the variation between the 𝐾 𝐷 𝐹𝑒 2+ -𝑀𝑔 of the cores and the rims of opx in the andesite and dacite using the equation ( 3), as in another study 18 :

∆ 𝑲 𝑫 𝑭𝒆 𝟐+ -𝑴𝒈 = ( [𝑋 𝑀𝑔𝑂 𝑙𝑖𝑞𝑢𝑖𝑑 𝑋 𝐹𝑒𝑂 𝑜𝑝𝑥 ] [𝑋 𝐹𝑒𝑂 𝑙𝑖𝑞𝑢𝑖𝑑 𝑋 𝑀𝑔𝑂 𝑜𝑝𝑥 ] ) 𝑐𝑜𝑟𝑒 -( [𝑋 𝑀𝑔𝑂 𝑙𝑖𝑞𝑢𝑖𝑑 𝑋 𝐹𝑒𝑂 𝑜𝑝𝑥 ] [𝑋 𝐹𝑒𝑂 𝑙𝑖𝑞𝑢𝑖𝑑 𝑋 𝑀𝑔𝑂 𝑜𝑝𝑥 ] ) 𝑟𝑖𝑚 (3)
With 𝑋 𝑗 𝑖 being the mole fraction of component j in phase i. For the liquid components, we took then mean compositions of the melt inclusions hosted in opx, studied in the andesite (5 MI) and the dacite (1 MI), with the ferric-ferrous ratio that was calculated using a calibrated equation from a study 122 and by knowing the temperature and fO2 (mgt/ilm pairs; Supplementary Data 3). Then, the variation in H2O content was deduced from equation ( 2), knowing the ∆ 𝑲 𝑫 𝑭𝒆 𝟐+ -𝑴𝒈 between the cores and the rims of opx.

CHAPTER VI.2: UNRAVELLING THE PRE-ERUPTIVE SPATIOTEMPORAL MAGMA DYNAMICS USING A MULTI-CRYSTAL APPROACH

VI.2.2.1. Geological context

The Kamchatka Volcanic Arc results from the subduction of the northeast corner of the Pacific plate (80-90 Ma) at a rate of 8-9 cm/yr beneath the Okhotsk plate (which is connected to the North American plate) (Figure VI.2.1a) (Gorbatov et al., 1997;[START_REF] Kayzar | Deciphering petrogenic processes using Pb isotope ratios from time-series samples at Bezymianny and Klyuchevskoy volcanoes, Central Kamchatka Depression[END_REF][START_REF] Konstantinovskaia | Arc-continent collision and subduction reversal in the Cenozoic evolution of the Northwest Pacific: an example from Kamchatka (NE Russia)[END_REF]Minster and Jordan, 1978). This arc has more than 25 active volcanic centers which are divided in three volcanic zones: the Eastern Volcanic Front (EVF), the Sredinny Ridge (SR) and the Central Kamchakta Depression (CKD), which is a graben due to intra-arc deformations because of the subduction (Churikova et al., 2001;[START_REF] Kayzar | Deciphering petrogenic processes using Pb isotope ratios from time-series samples at Bezymianny and Klyuchevskoy volcanoes, Central Kamchatka Depression[END_REF] and where the magma production is the largest [START_REF] Park | A Dangling Slab, Amplified Arc Volcanism, Mantle Flow and Seismic Anisotropy in the Kamchatka Plate Corner[END_REF][START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF][START_REF] Shapiro | Deep and shallow long-period volcanic seismicity linked by fluid-pressure transfer[END_REF]. The first signs of reactivation of the volcano were observed on 16 October 2010 with the formation of two fumarole fields at the top of the snow-covered Kizimen (Auer et al., 2018;[START_REF] Malik | The eruption of Kizimen volcano from October 2010 to March 2011[END_REF][START_REF] Melnikov | The 2010-2011 eruption of Kizimen volcano, Kamchatka: dynamics of eruptive activity and geologic-geomorphological impact (based on remote sensing data)[END_REF][START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF]. Then, on 11 November 2010, ash emissions up to 1 km in altitude began at the fumarole areas, marking the beginning of the eruption, as well as explosive episodes that generated pyroclastic density currents (Auer et al., 2018). Three main phases of the eruption have been identified by remote sensing techniques, seismic and visual observations (Firstov and Shakirova, 2014). The first one began on 9 December 2010 up to 11 May 2011 and resulted in the destruction of the upper part of the conduit and a summit extrusion with strong explosions and eruptions of pyroclastic density currents (Firstov and Shakirova, 2014). Then, between May and October 2011, a second phase involved extrusion of a lava flow down the eastern slope towards of the volcano. Finally, a third phase (October 2011 to June 2012) was outlined by the formation of a second branch of the lava flow which separated from the first one at 1300 m (Firstov and Shakirova, 2014). Frequent emission of pyroclastic density currents due to collapses of the lava dome and the front of the lava flow happened between 25 November 2011 to 12 January 2012 (Firstov and Shakirova, 2014). Between June 2012 and January 2013, slow squeezing-out of an extrusive dome with low explosive activity was observed (Firstov and Shakirova, 2014). A slow extrusion continued until December 2013, when the seismic tremor disappears below the detection limit (Auer et al., 2018).

For this eruption, a model of the pre-eruptive dynamics proposed that the emitted dacitic magma originated from the differentiation of a basaltic magma which was not erupted at the surface (Auer et al., 2018). Several models have assessed the presence of a stable silicic reservoir at 5-11 km under Kizimen, periodically refilled by magmatic injections (Auer et al., 2018;Churikova et al., 2013bChurikova et al., , 2007)).

Another seismic zone located during the eruption, from the ground surface to ~5 km, probably corresponds to the conduit along which silicic magma ascended to the surface (Auer et al., 2018). A mafic recharge is supposed to be responsible of the complete magma mixing of the lower part of the dacitic reservoir, forming an andesitic magma at the base of the magmatic reservoir (Auer et al., 2018). Such injection would have possibly occurred in 1963, simultaneously with a seismic crisis [START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF]. A large portion of the dacitic magma was not affected by this mafic recharge, leading to the formation of a zoned reservoir. Following opx-liquid thermometer calculations in the andesites, a temperature of 875°C is found, with an average pressure of 290 MPa (13 km for a density of 2300 kg/m 3 ) (Auer et al., 2018;Putirka, 2008). Prior to the 2010 eruption, an inhomogeneous mixing event occurred shortly between the andesitic magma and the dacitic magma due to some degassing of this latter, with incomplete mixing between the two magmas (Auer et al., 2018). This model was recently refined by a study focused on orthopyroxenes (opx), present both in the dacitic and the andesitic magmas (Ostorero et al., accepted). Textures, zonations as well as timescales modelling was carried out. 30 % of opx are zoned. In the dacite, the majority of opx are normal-zoned whereas, in the andesite, reverse-zoned opx are more represented. Opx compositions are spread over an En44-73 compositional range (Enstatite content (𝐸𝑛 = 𝑀𝑔 𝑀𝑔+𝐹𝑒

)). According to the opx zonations and pre-eruptive timescales modelling, magma mixing occurred between the andesitic and dacitic part of the reservoir 1.5 years period prior to the eruption (80 % of the timescales determined for this process whatever the type of zoning (normal or reverse)), while the upper part of the reservoir cooled down (Ostorero et al., accepted). These timescales are correlated to the seismic crisis that occurred 18 months before the eruption (Firstov and Shakirova, 2014;Ostorero et al., accepted;[START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF][START_REF] Shakirova | Features of the Kizimen Volcano area seismicity prior to and during the 2010-2013 eruption[END_REF].

VI.2.3. Material and methods

VI.2.3.1. Sampling in the field: andesites, dacites and banded samples

The samples were collected in the pyroclastic deposits of the 2010-2013 eruption, which emitted dacites during a three-week field mission which took place in August 2019 (

VI.2.3.2. Sample preparation

The unaltered dacitic and andesitic samples were crushed (up to 3 mm). One half was then crushed again into fine powder for whole rock analysis while the other half was sieved into different size fractions (1 mm-710 µm, 750-500 µm, 500-355 µm, 355-250 µm, 250-125 µm). These fractions have been washed in ultrasonic bath and dried at 80°C for 48h. They were then observed under the binocular microscope to select the fractions in which the crystals were the more abundant and automorphic. Orthopyroxenes (opx), plagioclase (plag), Fe-Ti oxides (magnetites, mgt) and amphiboles (amph) were then handpicked under the binocular microscope from size fractions ranging from 125 µm to 710 µm. They were aligned and mounted in epoxy resin and polished up to 0.3 µm to the middle part of the crystals for textural observations and analyses.

1832 opx (471 for the dacite and 1361 for the andesite) were studied for the eruption. Opx were oriented with c-axis in a north-south direction as they have been used for intracrystalline diffusion modelling (Ostorero et al., accepted). 1127 plag were studied in the fractions 125 to 710 µm (443 for the dacite and 944 for the andesite). 1175 mgt were separated in the fractions 250-355 µm and 125-250 µm (305 for the dacite and 870 for the andesite).

VI.2.3.3. Textural observations: Scanning Electron Microscope (SEM)

All the crystals (plag, opx, mgt) were observed by using Scanning Electron Microscope (SEM), Zeiss Supra 55VP (Sorbonne Université, Institut des Sciences de la Terre de Paris, ISTeP, Paris) and the Carl Zeiss EVO MA10 SEM at the PARI platform at the Institut de physique du globe de Paris (Université Paris Cité) using an acceleration voltage of 20 kV and a beam current of 8 nA. Contrast and brightness adjustments were made during each session to suit the crystal types and sizes.

Opx: The method used to identify the unzoned and zoned crystals (zonations have grey levels of different intensities) was the same as in Ostorero et al. (accepted) (Chapter VI.1).

Plag:

The zonations are characterized by the variations of grey levels of different intensities, reflecting the variations in composition in major elements as Ca. The same parameters as for opx are used. Highresolution image (mag x240) were also taken to have a better view of the textures and zonations (N = 7, 7 integrations of each line). Moreover, some chemical maps have been obtained to better apprehend the repartitions of different elements: Na, Ca, K, Al, Si, Mg, Sr and Ti and to highlight the different zonations.

Mgt: The possible chemical zonations (variations in Ti content) are recognized by the presence of a melt inclusions ring (Boudon et al., 2015). Images allow us to position the rim-core profiles to be carried out with the Electron Probe Micro-Analysis.

VI.2.3.4. Compositional analysis of crystals by Electron Probe Micro-Analysis (EPMA)

Opx: The method is detailed in Ostorero et al. (accepted). Crystals have been analysed for major elements (Si, Al, Ca, Mg, Na, K, Ti, Fe, Mn et P) with an acceleration voltage of 15 kV, a beam current of 10 nA and a focused beam of 2 µm (CAMECA SX-Five and SX-100; Service Camparis, Paris). Counting time on peak and background for Fe and Mg was set at 80 s and at 10 s for the other elements. The core to rim compositional profiles in zoned opx had a 2 µm step and an average length of 100 µm (~4 hours per profile). They were acquired perpendicular to the long axes of the opx and away from the corners to avoid three-dimensional effects such as growth (Allan et al., 2013;Fabbro et al., 2018;Krimer and Costa, 2017;[START_REF] Ostorero | Time -window into the transcrustal plumbing system dynamics of Dominica (Lesser Antilles)[END_REF].

Plag: Rim to core profiles are analyzed for major elements Ca, Na, Al, Si, O, K, Fe, as well as for the minor and trace elements Mg, Ti, Sr, Ba. For this purpose, two different programs have been prepared. The first one is used to measure the major elements using a voltage of 15 kV and a beam intensity of 10 nA with a beam diameter of 3 µm. The acquisition time is of 60 s and the step between each point is of 5 µm. The second program allows to measure trace elements with a voltage of 15 kV but a beam intensity of 300 nA. The spacing between the points and the size of the beam does not change compared to the previous program. To decrease the detection limits, the counting time is extended to 120 s. The automorphic plagioclase crystals that had well-delimited and preserved rims were analysed (to avoid growth effects). Furthermore, crystals grow mainly in one direction, neglecting the growth of a crystal tends to overestimate the timescales. However, the profiles have all been made along the b-or a-axis (perpendicular to the c-axis) where the effect of growth is negligible. It is also necessary to analyse enough plagioclases so that the compositions are statistically representative.

Mgt: The same analytical conditions as opx were used (15 kV, 10 nA and a focused beam of 2 µm) and a counting time of 10s was set for all elements except for Fe, Ti and Al (60 s).

VI.2.3.5. Magma dynamics through the Crystal System Analysis (CSA)

The crystals present in volcanic rocks grow from a magmatic liquid in which they reside by incorporation of the chemical elements constituting them. This incorporation is governed by thermodynamic and kinetic laws. The crystals are thus evolving between different magmatic environments (ME) defined as the conditions of pressure, temperature, oxygen fugacity, fluid partial pressure, and chemical composition of the liquid and thus may help us to track the pre-eruptive dynamics of the plumbing system [START_REF] Anderson | Probable relations between plagioclase zoning and magma dynamics, Fuego Volcano, Guatemala[END_REF][START_REF] Giacomoni | Plagioclase as archive of magma ascent dynamics on "open conduit" volcanoes: The 2001-2006 eruptive period at[END_REF]Kahl et al., 2011;Petrone and Mangler, 2021;[START_REF] Streck | Mineral Textures and Zoning as Evidence for Open System Processes[END_REF]. The zoning patterns often consist of two or more compositional plateaus (referred as "rim") with diffusive boundaries separating them (Kahl et al., 2015). Studying the evolution of the composition of a crystal from the core to the rims then allows to know the chemical history of a particular crystal. Applying statistically this compositional study on several crystals permits to investigate in time and space the changes of crystallization conditions.

Classically, the ME are represented by boxes, the size of which is proportional to the ME contribution, and arrows between them indicate the passage, from the core to the rim, of the recorded environments for single-zoned or multiple-zoned crystals (Kahl et al., 2011;[START_REF] Ostorero | Time -window into the transcrustal plumbing system dynamics of Dominica (Lesser Antilles)[END_REF][START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF].

For opx: The En content may increase (reverse-zoned opx), which is linked to a temperature (T) increase, with or without, magma mixing with a hotter and more basic magma (Martel et al., 1999;[START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF]. Reverse-zoned rims can also be created following a decrease of water content in the melt from degassing processes (Waters and Lange, 2017). Conversely, a decrease in the En content (normal-zoned opx) may be linked to a pressure drop, decrease in melt H2O content or mixing with a colder magma [START_REF] Kahl | Compositionally zoned crystals and real-time degassing data reveal changes in magma transfer dynamics during the 2006 summit eruptive episodes of Mt. Etna[END_REF]Martel et al., 1999;Saunders et al., 2012;[START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF] or a higher water content in the melt (Waters and Lange, 2017).

For plag: The An content may increase (reverse-zoned plag) following an increase of T [START_REF] Ustunisik | New insights into the processes controlling compositional zoning in plagioclase[END_REF]. Reverse zonation can also be caused by convection in a polybaric chamber but also by magma decompression, e.g. when magma ascends to a shallower chamber while An content may decrease, indicating a decrease in temperature of the reservoir [START_REF] Ustunisik | New insights into the processes controlling compositional zoning in plagioclase[END_REF]. Waters and Lange (2017) also evidenced that normal-zoned plagioclase (sodic rims) generally accompany reversely zoned opx in the case of loss of H2O during degassing.

For Mgt: If magnetites are unexsolved, Ti content may be used to constrain T and/or fO2 changes during the late stages of magma ascent in the conduit and magnetite crystallization, shortly before quenching (timescales of the order of days to months) (Boudon et al., 2015;Nakamura, 1995a). Ti profiles are linked to T variations, with reverse zonation produced by a T increase and a normal zonation by a T decrease [START_REF] Devine | Magma Storage Region Processes Inferred from Geochemistry of Fe-Ti Oxides in Andesitic Magma, Soufriere Hills Volcano, Montserrat[END_REF].

VI.2.3.6. Timescales of magma dynamics

During a volcanic eruption, the drop in temperature freezes the element diffusion and allows to keep the memory of the chemical exchange if the equilibrium has not been reached. Working backwards from the known date of eruption, transfer timescales between one ME to another can be determined. It corresponds to the time between the disrupting event at depth or in the conduit, at the origin of the change of composition within the crystal, and the moment when the magma comes out and is frozen at the surface. The diffusion process is described by the two thermodynamic Fick's laws which relate the diffusion coefficient to the concentration gradient of the studied element, as well as the variation of concentration in time. The diffusion coefficient depends theoretically on parameters such as chemical composition (Xi; molar fraction of the mineral constituent element), crystallographic axis, temperature (T in Kelvin), pressure (P in Pa), oxygen fugacity (fO2) and water fugacity (fH2O) (Costa and Morgan., 2010;Costa et al., 2020). The diffusion rate is strongly dependent on the temperature which is the parameter that has the strongest influence on the calculated diffusion timescales. The calculation of the diffusion timescales is based on the diffusion coefficient of the element in the crystal conditions and its initial profile. The equation of the diffusion coefficient follows an Arrhenius equation (VI. 2.1):

𝐷 = 𝐷 0 𝑒 -𝑄 𝑅𝑇 (VI. 2.1)
Where D is the element diffusion coefficient (m². s -1 ), 𝐷 0 is a constant, Q is the activation energy (J.mol -1 ), R is the molar gas constant (8.31 J. K -1 .mol -1 ) and T is the temperature in Kelvin.

VI.2.3.7. Diffusion timescales models

Opx: Modelling of the timescales associated to the zonations in the zoned crystals has been done by following the method developed in Allan et al. (2013) and Couperthwaite et al. (2020), using the parametrization of the Fe-Mg interdiffusion coefficient D (Ganguly and Tazzoli, 1994). Here we only consider the T effect with a diffusion coefficient that cover the range of opx compositions studied here. Opx textures and compositions from Ostorero et al. (accepted) are summarized partly in the Results, as they are compared to plag and mgt and as they will be used in the Crystal System Analysis but the timescales modelled in the previous study will not be described in details.

Plag: Trace elements such as magnesium (Mg) diffuse fast enough (7.1×10 -8 m²/s along the b-axis) to obtain a diffusion profile (Druitt et al., 2012). The diffusion coefficient of the Mg element in plagioclase is given by the equation (VI. 2.2) (LaTourrette and Wasserburg, 1998): 92 . 10 -4.1𝑋 𝐴𝑛 -3.1 e -266 000 𝑅𝑇 (VI. 2.2) Where 𝐷 𝑝𝑙𝑎𝑔 𝑀𝑔 is in m². s -1 , 𝑋 𝐴𝑛 is the mole fraction of An and T is in Kelvin.

𝐷 𝑝𝑙𝑎𝑔 𝑀𝑔 = 2.
As the interdiffusion of the major elements CaAl-NaSi is very slow, one can consider that the profile in An is preserved, which gives access to the initial conditions. The temperatures used for the timescales modeling are those determined in magnetite/ilmenite pairs of the same samples in another study (847 ± 57°C for the andesitic host melt and 832 ± 37°C for the dacitic one) (Ostorero et al., accepted;Sauerzapf et al., 2008).

The input parameters are:

-𝑋 𝐴𝑛 , the mole fraction of An in the profile, calculated from the major elements of the EPMA analyses; -The Mg profile measured with the EPMA (ppm); -The initial Mg profile is the initial magnesium contents, i. ) (Druitt et al., 2012) between the magma and the plagioclases to recover the initial content [START_REF] Bindeman | Ion Microprobe Study of Plagioclase-Basalt Partition Experiments at Natural Concentration Levels of Trace Elements[END_REF]) (equation VI.

3h).

The only two types present in the dacite are plag with normal zonations (53% of the population, type 2; Table VI.2.1) and plag with oscillatory zonations (47% of the population, type 4; Table VI.2.1). In the dacitic sample, no plag with RZ (type 1) or plag with a "patched" core (type 3) were found. In andesite, the majority of plag (75% of type 2 and 3) show NZ, out of which 43% have a patched core (only type 3) (Table VI.2.1). The remaining 25% show a RZ at the rim with a resorption surface (type 1) (Table VI.2.1). Oscillatory zonation (type 4) is not observed in the andesitic samples (Table VI.2 Mgt: Out of 1175 mgt, 16 profiles have been analysed in the samples (5 for the dacite and 11 for the andesite, Supplementary Data VI.2.2). Only 3 profiles in the dacite were suitable for the modelling of the timescales due to bumps in the profiles, contrary to the andesite were all transect were modelled. In the dacite, one timescale was modelled in a normal-zoned mgt and two timescales have been modelled in multiple-zoned mgt, from the inner to the outer rim. All the other profiles were modelled in singlezoned mgt (Supplementary Data VI.

VI.2.5. Discussion

VI.2.5.1. Dynamics of the magma reservoir: a Crystal System Analysis (CSA) approach

We applied the CSA to study the links between the magmatic environment (ME) recorded by the crystals through time, before the eruption (Kahl et al., 2017(Kahl et al., , 2015[START_REF] Kahl | Compositionally zoned crystals and real-time degassing data reveal changes in magma transfer dynamics during the 2006 summit eruptive episodes of Mt. Etna[END_REF](Kahl et al., , 2011)). The ME can be seen as virtual areas in a zoned magma reservoir or different magma batches stored in the mush (Kahl et al., 2017).

Cooling or differentiation would lead to continuous variations of compositions whereas the connections between the ME can be interpreted as evidences of magmatic processes taking place in magmatic reservoirs or during magma ascent. The crystals can either move from one ME to another one (injection of magma, crystal settling…) or the environment can change around them (due to volatile transfer, degassing) which results in changes of the crystallization conditions [START_REF] Bachmann | On the Origin of Crystal-poor Rhyolites: Extracted from Batholithic Crystal Mushes[END_REF]Bergantz, 2006, 2003;Caricchi and Blundy, 2015;Huber et al., 2010;Pistone et al., 2013) VI.2.5.1.1. Cores: evidence of an initial dacitic magma (ME1)

The compositions of the emitted plag cores describe a wide range of compositions, An42-54 (Figure VI.2.4b; Figure VI.2.8a; c). in both dacitic and andesitic magmas that define a first magmatic environment (ME1). As plag, all the opx, unzoned and single-zoned ones, in the dacite and the andesite, display the same core compositions. This was interpreted as evidence of the common origin of all opx and plag from the initial dacitic magma (Ostorero et al., accepted). This ME thus also corresponds to the unzoned opx and cores compositions of the zoned opx of the dacite and andesite (En62-64; 4d and6c). This characterizes a second magmatic environment (ME2), which is only recorded in opx by 2 % of the zoned crystals (Ostorero et al., accepted). All of these crystals pass from ME1 with En62-64 and An42-54 to ME2 defined by An70-84. 43 % of these plag show an important partial (patched core) or total core resorption (Figure VI.2.3-8c). The resorption and the sudden jumps of composition between the core and the intermediate zone can be interpreted as an evidence of the injection into the initial dacitic reservoir of a hotter magma of more mafic composition. This results in a mixing process (mixing 1) creating the andesitic magma. However, 25 % of plag have an An-poor intermediate zone (An58-46) (Figure VI. 2.4c); this suggests that these crystals were not modified by the injection and stayed in ME1(Figure VI.2.8c). The mixing event does not involve all the dacitic magma; explaining why the An-rich melted zone is not observed in the crystals of the dacitic samples or in the opx (even if a small proportions of them have resorption zones; (Ostorero et al., accepted)) (Figure VI.2.8a). VI.2.5.1.3. The rims: witnesses of different mixing processes in a zoned magma reservoir After the first mixing event, the magma reservoir is zoned, with a dacitic magma at the top and an andesitic magma at the base. In andesite, all plag rims are between An54 and An70 ( This may be interpreted as a temperature difference between ME1 and ME3, with ME3 being hotter than ME1 (Kahl et al., 2011;Martel et al., 1999;[START_REF] Ustunisik | New insights into the processes controlling compositional zoning in plagioclase[END_REF]. As seen previously, the temperature of dacitic magma (ME1) is estimated at about 832 °C while the temperature of the andesitic magma is higher, around 847 °C (ME3). This argues in favor of a second mixing process at the boundary between dacitic and andesitic magma in the temperature-zoned magma reservoir. An additional argument in favor of such a mixing process is brought by the signature in trace elements (Fe) of the different plag populations (Figure VI.2.5b). The mixing of dacitic magma (Fe-poor) with andesitic magma (Fe-rich) is responsible for the Fe peak in ME3 rim of plag coming from the dacite (Figure

VI.2.5b).

The plag with An-rich intermediate zones that were already in the andesitic magma are cooled down due to the mixing with the dacite and crystallize the same ME3 rim as plag coming from the dacite.

The majority of the opx are unzoned (70%) (Ostorero et al.,accepted) (Figure VI.2.8b;d). One hypothesis to explain this high proportion of unzoned opx is that these crystals were disturbed by the formation of the andesitic magma but their compositions were homogenized by re-equilibration of the opx composition with time. But the aphyric basalt must have been hotter, leading to mafic En contents in the opx rims but as there are no witnesses in the opx cores or unzoned opx of these more mafic homogenized compositions (En content > En 62-64), this hypothesis is thus less likely.

Another hypothesis could be that, as only 30% of the opx crystals experienced the changes of the magmatic conditions, it would mean that a small part of the reservoir was affected by these changes, due to the high viscosity of the dacitic magma (Auer et al., 2018;Ostorero et al., accepted;Ubide et al., 2014) (Figure VI.2.6). In the andesite, reverse-zoned opx rims are produced (change in composition from ME1 to ME3), as well as a few normal-zoned opx (ME1 to ME4). Multiple-zoned opx describe the three ME (ME1; 2 and 3) (Figure VI.2.8b; d). As plag, the single-zoned opx thus also recorded the mixing process at the interface between dacite and andesite: opx from the dacite are heated when incorporated in andesite melt, whereas the opx from the andesite are cooled down by the dacitic magma. For the dacite, the major changes recorded in opx are between ME1 to ME4 (normal zoning; with a small proportion of opx that also record reverse zoning, with a change from ME1 to ME3), indicating some cooling ( Looking at plag, we have first looked at the processes responsible for significant changes in compositions in An. However, there are also An fluctuations at micrometer scale (~5 µm) creating oscillatory zonations. These peaks are related to the changing conditions that can be caused by magmatic convection within the reservoir [START_REF] Ustunisik | New insights into the processes controlling compositional zoning in plagioclase[END_REF]. They are particularly observable in the plag of type 4 in dacites (47%). The cores of plag in andesite frequently show traces of convection (large ranges of An content; (Figure VI.2.4b). To explain the oscillatory zonations of the plag in the dacite, it could also result from localized magma movements within the zoned reservoir (Couch et al., 2001). The hypothesis of convection is also supported by opx in the andesite (22%) and dacite (5%), mostly in ME1 (zonations in the same En ranges, En62-64) (Figure VI.2.8b;d) (Ostorero et al., accepted). These fluctuations may therefore result from exchanges between the different ME but also of local variations in T, P and compositions. This type of zonation reinforces the idea of the dynamical character of the reservoir and the frequent exchanges between the two magmas during the years preceding the eruption. By comparing plag and opx composition, we thus identified two main mixing processes in the magma reservoir history.

VI.2.5.2. Timescales of pre-eruptive magmatic processes

The first mixing event leads to the formation of a zoned magma reservoir by mixing of the resident dacitic magma with an aphyric basic magma, as recorded by the intermediate zones and rims of some plag in the andesite (Figure VI. 2.8c). This timescale of this event cannot be precisely constrained with diffusion because the injection melted the cores of plag and the temporal information disappeared. But, from the other timescales in plag, they suggest that the growth of the rims that followed the resorption event took place in the 5 years before the eruption. So, the injection occurred more than 5 years prior eruption. Considering the time corresponding to the resorption of the cores as well as the crystallization time of the intermediate zone (~100 µm), this recharge may have taken place in the few decades preceding the eruption.

The second mixing process occurred at the interface between the remnant dacitic magma and the andesitic magma and was clearly recorded by opx from 1.5 1.5 years before the eruption, the zoned opx and plag record a progressive heating event in the andesitic part and a cooling process in the dacite. These events are more dynamic towards the onset of the eruption, as 60% of the timescales of opx are below 1 year before the eruption and with the short timescales that are recorded by the plag. Magma mixing between the dacite and the andesite is then efficiently increasing. In the last two weeks before the eruption, mgt either begin to record this magma mixing process or magma ascent in the conduit. This magmatic heating process is correlated to the onset of the seismic crisis that lasted 1.5 years at Kizimen (Auer et al., 2018;[START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF].

By looking at the cumulative number of timescales (Figure VI.2.9a-c), we can see that the timescale curve associated to the reverse-zoned plag has a very steep slope, meaning that the mixing process creating the andesitic magma was very dynamic. The opx timescales give more progressive timescales, with a high majority below 1.5 years before the eruption and an increase of the timescales 1 year before the eruption, meaning that the system was more dynamic, at the time of the mixing between dacite and andesite (Figure VI.2.9a-c).

Furthermore, Ti diffusion modelling in mgt is used to study the transit time of magma in the conduit or early phases of pre-eruptive heating as their compositions change much faster than silicates in response to magma system temperature rises (Boudon et al., 2015;[START_REF] Devine | Magma Storage Region Processes Inferred from Geochemistry of Fe-Ti Oxides in Andesitic Magma, Soufriere Hills Volcano, Montserrat[END_REF][START_REF] Morgado | Localised heating and intensive magmatic conditions prior to the 22-23 April 2015 Calbuco volcano eruption (Southern Chile)[END_REF][START_REF] Venezky | Petrology and Fe-Ti oxide reequilibration of the 1991 Mount Unzen mixed magma[END_REF]. Here, the timescales of mgt of the two parts of the reservoir (either dacitic or andesitic) are very young (< 15 days). These timescales can then be linked to a heating event in the andesite, as the mgt are mostly reverse-zoned and a cooling event in the dacite, as evidenced by opx and plag in the dacite (even if a lesser number of timescales are studied) [START_REF] Devine | Magma Storage Region Processes Inferred from Geochemistry of Fe-Ti Oxides in Andesitic Magma, Soufriere Hills Volcano, Montserrat[END_REF]. However, it is known from geothermometry that Fe-Ti oxides will be open to diffusion to lower temperatures than silicates [START_REF] Frost | Fe-Ti oxide-silicate equilibria; assemblages with fayalitic olivine[END_REF][START_REF] Frost | Preservation of Fe isotope heterogeneities during diagenesis and metamorphism of banded iron formation[END_REF]. For Mg in plag, and 𝑋 𝐴𝑛 = 0.67 (which is in the An range of the plag of this study), the closure temperatures at which diffusion stops was estimated between 601-872 °C (Costa, 2021;[START_REF] Van Orman | Magnesium diffusion in plagioclase: Dependence on composition, and implications for thermal resetting of the 26 Al-26 Mg early solar system chronometer[END_REF] and for Fe-Mg interdiffusion in opx, for high En opx (En91 and En98), they were estimated to vary between ~800-1100 °C (Costa, 2021;Dohmen et al., 2016). We can thus hypothesize that, as the closure temperatures of mgt are lower than plag and opx, the timescales modelled in the zoned mgt are not necessarily corresponding to magmatic processes that occurred in the reservoir. They could also be related to magma ascent in the conduit: as after the explosive phase, a lava dome was growing and emitted a lava flow. The Ti diffusion in mgt may have been frozen in such lava flow, as it occurred for Montagne Pelée in the order of several months (Boudon et al., 2015). According to these results, given the types 2 and 3 timescales in plag, the recharge event responsible of the zonation of the reservoir occurred more than 5 years prior eruption (i.e. necessarily before 2005) but was not constrained with plag, mgt and opx. But, as an intense seismic crisis occurred in 1963, the injection could have occurred around this time (Auer et al., 2018;[START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF]. About the beginning of mixing recorded by plag from 1-5 years, 20 % of opx timescales record some magmatic process from 2003 to 2008. This event could have been detected by a monitoring network with more than one seismic station, as the timescale curve shape of plag is similar to that related to the vigorous mixing between dacite and andesite that is correlated with seismic record in 2009 (Figure VI.2.9a).

By looking at the main minerals in the emitted products of the 2010-2013 eruption of Kizimen, we thus propose a revised model of the pre-eruptive processes and their chronology. ). From 1.5 years and during the months preceding the eruption, the mixing between the dacitic and andesitic reservoir becomes more intense, in correlation with the seismicity (Auer et al., 2018;[START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF]. The incorporated plag resorb partially before crystallizing a new rim of composition An60-70 (Figure VI.2.10e). These processes in the reservoir may be linked to the increasing seismic activity.

In November 2010, a third magma mixing event involving mechanical mixing without homogenization (mingling), occurred and probably triggered magma ascent to the surface carrying both types of mingled magma and producing the banded lava observed on the field (Figure VI.2.10f).

VI.2.6. Conclusion

Thus, interactions between four magmatic environments have been highlighted for the crystal assemblage of Kizimen 2010 eruption (plag, opx, mgt), with different proportions recording these An changes and En changes. Considering these three types of minerals and their pre-eruptive timescales, the history of the magma plumbing system can be refined. Magma dynamics highlight an injection event followed by a first mixing event and then a second magma mixing process mainly 1.5 years prior eruption. A third incomplete magma mixing event with mingling textures and eruption was probably followed by magma ascent in the lava dome. Adding the amphiboles, it could lead to a complete model of the pre-eruptive magma dynamics prior to the eruption of 2010 of Kizimen.

VI.3.1. Introduction

Nowadays, an important debate in volcanology is the triggering of eruptions and the mechanisms leading to large explosive eruptions (Arzilli et al., 2019b;Caricchi et al., 2014a;[START_REF] Gregg | Catastrophic caldera-forming eruptions II: The subordinate role of magma buoyancy as an eruption trigger[END_REF][START_REF] Gudmundsson | The mechanics of large volcanic eruptions[END_REF][START_REF] Malfait | Supervolcano eruptions driven by melt buoyancy in large silicic magma chambers[END_REF]. On the one hand, simulations and densities measurements show that magma buoyancy (due to the density contrast between the hot partially molten magma and the surrounding rock), potentially producing overpressures in large silicic systems, is key to triggering super-eruptions whereas pressurization associated with a magma injection can lead to relatively small and frequent eruptions (Caricchi et al., 2014a;[START_REF] Malfait | Supervolcano eruptions driven by melt buoyancy in large silicic magma chambers[END_REF]. However, other studies showed that magma buoyancy does not necessarily result in over pressurization of the reservoir [START_REF] Gregg | Catastrophic caldera-forming eruptions II: The subordinate role of magma buoyancy as an eruption trigger[END_REF].

These triggering mechanisms are difficult to constrain and can be explained by five main phenomenon (Arzilli et al., 2019b). The most commonly identified is the injection of volatile-rich magmas into a more evolved magma chamber [START_REF] Blake | Volcanism and the dynamics of open magma chambers[END_REF][START_REF] Cassidy | Volatile dilution during magma injections and implications for volcano explosivity[END_REF][START_REF] De Silva | Triggering explosive eruptions-The case for silicic magma recharge at Huaynaputina, southern Peru[END_REF][START_REF] Pallister | A basalt trigger for the 1991 eruptions of Pinatubo volcano?[END_REF][START_REF] Sigmundsson | Intrusion triggering of the 2010 Eyjafjallajökull explosive eruption[END_REF][START_REF] Walker | Tarawera 1886, New Zealand -A basaltic plinian fissure eruption[END_REF][START_REF] Wehrmann | The 12.4 ka Upper Apoyeque Tephra, Nicaragua: stratigraphy, dispersal, composition, magma reservoir conditions and trigger of the plinian eruption[END_REF][START_REF] Williams | The October 1902 plinian eruption of Santa Maria volcano, Guatemala[END_REF]), but others are identified, such as magma mixing following mafic magma injection [START_REF] Arienzo | Processes and timescales of magma evolution prior to the Campanian Ignimbrite eruption (Campi Flegrei, Italy)[END_REF][START_REF] Bardintzeff | Magma mixing processes in volcanic contexts, a thermodynamic approach with the examples of St Vincent Soufrière volcano, West lndies and Cerro Chiquito, Guatemala[END_REF]Chamberlain et al., 2014;[START_REF] Jellinek | A model for the origin of large silicic magma chambers: precursors of caldera-forming eruptions[END_REF][START_REF] Leonard | Basalt triggering of the c. AD 1305 Kaharoa rhyolite eruption, Tarawera Volcanic Complex, New Zealand[END_REF][START_REF] Macías | Storage conditions and magma processes triggering the 1818 CE Plinian eruption of Volcán de Colima[END_REF]Matthews et al., 2012;Morgavi et al., 2017[START_REF] Morgavi | The Grizzly Lake complex (Yellowstone Volcano, USA): Mixing between basalt and rhyolite unraveled by microanalysis and X-ray microtomography[END_REF][START_REF] Murphy | The role of magma mixing in triggering the current eruption at the Soufriere Hills Volcano, Montserrat, West Indies[END_REF]Perugini and Poli, 2012;Saunders et al., 2012;[START_REF] Sparks | Magma mixing: a mechanism for triggering acid explosive eruptions[END_REF][START_REF] Stock | Triggering of major eruptions recorded by actively forming cumulates[END_REF][START_REF] Wark | Pre-eruption recharge of the Bishop magma system[END_REF], generation of overpressure due to the exsolution of volatiles via second boiling due to protracted crystallization [START_REF] Stock | Late-stage volatile saturation as a potential trigger for explosive volcanic eruptions[END_REF][START_REF] Tramontano | Internal triggering of volcanic eruptions: tracking overpressure regimes for giant magma bodies[END_REF], perturbation of magmatic reservoirs close to a critical state by tectonic earthquakes [START_REF] Linde | Triggering of volcanic eruptions[END_REF]Seropian et al., 2021), self-mixing (Couch et al., 2001) and finally, sector collapse of a volcanic edifice [START_REF] Belousov | Directed blasts and blast-generated pyroclastic density currents: a comparison of the Bezymianny 1956, Mount St Helens 1980, and Soufrière Hills, Montserrat 1997 eruptions and deposits[END_REF][START_REF] Lipman | The 1980 eruptions of Mount St[END_REF].

The injection of mafic magma into more evolved magma causes superheating of the evolved magma and vigorous convection, which can favor vesiculation and increase magma pressure, leading to eruptions [START_REF] Sparks | Magma mixing: a mechanism for triggering acid explosive eruptions[END_REF][START_REF] Stock | Triggering of major eruptions recorded by actively forming cumulates[END_REF]. Convection models and petrological data show that the magmas physically mix shortly on timescales of minutes to hours (to several weeks) before eruption and the mafic magma is quenched, leading to crystallization, exsolution and transfer of volatiles to the evolved magma [START_REF] Leonard | Basalt triggering of the c. AD 1305 Kaharoa rhyolite eruption, Tarawera Volcanic Complex, New Zealand[END_REF][START_REF] Sparks | Magma mixing: a mechanism for triggering acid explosive eruptions[END_REF]. Indeed, large-scale felsic-mafic magma interactions are recognized as eruption triggers for the destruction of the Las Cañadas magma chamber [START_REF] Stock | Triggering of major eruptions recorded by actively forming cumulates[END_REF], when these mafic injections are large or frequent, they likely generate explosive eruptions at Colima volcano (Mexico) [START_REF] Macías | Storage conditions and magma processes triggering the 1818 CE Plinian eruption of Volcán de Colima[END_REF].

Magma mixing is the chemical interaction between two or more magmas that produces a composition intermediate between the original end-members [START_REF] Bunsen | Ueber die Processe der vulkanischen Gesteinsbildungen Islands[END_REF][START_REF] Jarvis | Crystal and Volatile Controls on the Mixing and Mingling of Magmas[END_REF][START_REF] Wilcox | The Idea of Magma Mixing: History of a Struggle for Acceptance[END_REF]. Chemical mixing is occurring with the diffusion of different elements along spatial gradients in chemical potential and can lead to the hybridization and homogeneous products [START_REF] Adkins | Equilibrium Thermodynamics[END_REF][START_REF] Humphreys | Magma hybridisation and diffusive exchange recorded in heterogeneous glasses from Soufrière Hills Volcano, Montserrat[END_REF][START_REF] Jarvis | Crystal and Volatile Controls on the Mixing and Mingling of Magmas[END_REF]. By contrast, mingling is the physical interaction of the two magmas, such as through convective stirring [START_REF] Oldenburg | Dynamic mixing in magma bodies: Theory, simulations, and implications[END_REF] or chaotic advection (Morgavi et al., 2013b;[START_REF] Perugini | Analysis and numerical simulation of chaotic advection and chemical diffusion during magma mixing: petrological implications[END_REF]) leaving compositional heterogeneities [START_REF] Jarvis | Crystal and Volatile Controls on the Mixing and Mingling of Magmas[END_REF]. However, physical mixing alone never occurs as chemical diffusion will always happen and the degree of homogenization is depending on temperature and time spent after the onset of contact between the end-members [START_REF] Laumonier | Experimental simulation of magma mixing at high pressure[END_REF]. Mingling textures are those that have cooled fast enough to prevent chemical homogenization, regardless of the process involved (static or dynamic where homogenization is physically enhanced via folding and stretching which shorten diffusion distances) [START_REF] Laumonier | Experimental simulation of magma mixing at high pressure[END_REF]Perugini et al., 2010[START_REF] Perugini | Trace element mobility during magma mixing: Preliminary experimental results[END_REF][START_REF] Perugini | Diffusive fractionation of trace elements by chaotic mixing of magmas[END_REF][START_REF] Perugini | Chaotic advection, fractals and diffusion during mixing of magmas: evidence from lava flows[END_REF][START_REF] Perugini | Analysis and numerical simulation of chaotic advection and chemical diffusion during magma mixing: petrological implications[END_REF][START_REF] Petrelli | Transition to chaos and implications for time-scales of magma hybridization during mixing processes in magma chambers[END_REF]. The morphology and textures of mixing/mingling patterns depend of the time evolution of the magmatic systems and they can be relatively short in natural magmatic reservoirs, of the order of hours [START_REF] Fourmentraux | Crystal fractionation, magma step ascent, and syn-eruptive mingling: the Averno 2 eruption (Phlegraean Fields, Italy)[END_REF]Montagna et al., 2015;Perugini et al., 2010). Estimating the timescales between mixing processes and eruption is thus of high importance as mixing is at stake before many eruptions. However, not all magmas that mix lead to eruptions and not all eruptions are triggered by magma mixing [START_REF] Sparks | Magma mixing: a mechanism for triggering acid explosive eruptions[END_REF].

After 83 years of quiescence, Kizimen volcano (Kamchatka peninsula) produced its first historical magmatic eruption on 11 November 2010, after 18 months of slow-going seismic precursory process (Firstov and Shakirova, 2014) and a phreatic eruption in 1927 (Auer et al., 2018;Melekestsev et al., 1995)) (Figure VI. 3.1). The 2010 eruption of Kizimen represents an important case study because some mafic magma was injected in the dacitic reservoir, potentially in 1963, along with an important seismic crisis, but did not immediately trigger the eruption (Auer et al., 2018;Ostorero et al., accepted). A first magma mixing between the basalt and the dacite led to the formation of an andesite. Then, beginning 1.5 years before the 2010 eruption, unrest and second magma mixing processes gradually occurred at Kizimen, at the contact between the andesitic magma, up to the mingling event shortly prior to the eruption (Ostorero et al., accepted). From there on, we will call this event the third magma mixing process. This last event generated mingling textures, as banded lavas were emitted along some silicarich andesites and dacites during the eruption (Figure VI. 3.1c) (Auer et al., 2018).

Studying the first mixing event that followed the mafic injection but that did not directly lead to the eruption of Kizimen is fundamental to understand the pre-eruptive magmatic processes that followed. Natural evidence of magma mixing processes might serve as chronometers to estimate the time interval between mixing and eruption, as the mixing patterns depend on the time evolution of the magmatic systems [START_REF] Morgavi | Decrypting Magma Mixing in Igneous Systems[END_REF](Morgavi et al., , 2017(Morgavi et al., , 2013b(Morgavi et al., , 2013a;;Perugini et al., 2015aPerugini et al., , 2010;;Perugini and Poli, 2012). Performing magma mixing experiments and numerical simulations, short timescales are obtained for these events, as shown before two pyroclastic sequences in Campi Flegrei, Italy, where mixing processes lasted approximately 2 to 9 days, prior to eruption (Perugini et al., 2010) or a few hours to minutes (Montagna et al., 2015;Perugini et al., 2015a). Constraining such timescales is thus of high importance for volcanologists to mitigate volcanic hazards. Furthermore, magma mixing leads to timedependent diffusion of chemical elements. Phenocrysts textures and their sizes distribution also give insights into crystallization and differentiation processes and into the changes of storage conditions during processes such as magma mixing, cooling, degassing and decompression [START_REF] Hawkesworth | Time scales of magmatic processes[END_REF][START_REF] Jarvis | Crystal and Volatile Controls on the Mixing and Mingling of Magmas[END_REF]. Indeed, during magma ascent, decompression results in water exsolution and an increase of the liquidus temperature, resulting in the crystallization of small and out of equilibrium anhydrous minerals [START_REF] Hammer | An experimental study of the kinetics of decompression-induced crystallization in silicic melt[END_REF]. Cooling and crystallization of the new injected melt and associated latent heat release should control the dynamics of the system [START_REF] Jarvis | Crystal and Volatile Controls on the Mixing and Mingling of Magmas[END_REF].

Experiments that studied the kinetics and chemical aspects of mixing were previously performed [START_REF] De Campos | Enhancement of magma mixing efficiency by chaotic dynamics: an experimental study[END_REF], 2008, 2004;Kouchi andSunagawa, 1985, 1983;[START_REF] Laumonier | Experimental mixing of hydrous magmas[END_REF][START_REF] Laumonier | Experimental simulation of magma mixing at high pressure[END_REF]Morgavi et al., 2013b[START_REF] Morgavi | Decrypting Magma Mixing in Igneous Systems[END_REF]Morgavi et al., , 2017Morgavi et al., , 2013a;;Perugini et al., 2015aPerugini et al., , 2010;;Perugini and Poli, 2012), as well as crystallization experiments [START_REF] Arzilli | Plagioclase nucleation and growth kinetics in a hydrous basaltic melt by decompression experiments[END_REF][START_REF] Arzilli | Crystallization kinetics of alkali feldspars in cooling and decompression-induced crystallization experiments in trachytic melt[END_REF][START_REF] Erdmann | Textures of Peritectic Crystals as Guides to Reactive Minerals in Magmatic Systems: New Insights from Melting Experiments[END_REF][START_REF] Hammer | An experimental study of the kinetics of decompression-induced crystallization in silicic melt[END_REF][START_REF] Mollard | Decompression-induced Crystallization in Hydrated Silica-rich Melts: Empirical Models of Experimental Plagioclase Nucleation and Growth Kinetics[END_REF][START_REF] Shea | Kinetics of cooling-and decompression-induced crystallization in hydrous mafic-intermediate magmas[END_REF]. Mixing and crystallization experiments are thus missing to proper quantify the time interval between the mixing process and eruptions. The aim of this study is to better understand the timescales of magma mixing events by combining the study of the crystallization kinetics as well as mixing.

We focus on constraining the kinetics of magma mixing following a mafic injection inside the dacitic reservoir of Kizimen 50 years before the 2010 eruption of Kizimen, as no basalt was found on the field. This intrusion was thought to be of mafic origin, as traces of resorbed olivines are found in the eruptive products of the 2010 eruption (Auer et al., 2018). To do this, we performed for the first-time chemical mixing at sub-liquidus conditions via a chaotic mixing experiment using a natural dacite and a synthetic basalt.

The second event in which we are interested in is the third mixing event, shortly prior eruption (the mingling phase). As this third mixing event, that involved the dacitic magma and the andesitic magma, occurred in an unknown timeframe just before the eruption, it is of interest to understand the triggering event at Kizimen, by studying the banded lavas. Several types of timescales experienced by the magmatic reservoir are thus investigated here. This study of the mixing timescales, as well as natural banded lavas from this eruption can serve for comparison for assessing eruptive mechanisms and hazards on similar volcanoes such as Mt Unzen and Mt Sinabung, also sources of poorly mixed magmas (Auer et al., 2018).

VI.3.2. Material and methods

VI.3.2.1. Natural samples

VI.3.2.2. End-member selection and preparation

The experiment was performed using a natural dacitic and a synthetic basalt end-member (Supplementary Figure VI.3.1). The dacitic sample is porphyritic and consists of ~29% of crystals (plagioclase, amphiboles, orthopyroxenes, Fe-Ti oxides and some resorbed olivines and quartz) (Ostorero et al., accepted). The synthetic basalt composition is the equivalent of a basaltic enclave in a lava flow of Kizimen (Churikova et al., 2013b)

VI.3.2.3. High-temperature magma mixing experiment

The apparatus used for magma mixing experiments is the Chaotic Magma Mixing Apparatus (hereafter COMMA), which aims to mix viscous natural and synthetic melts up to 1500 °C (Figure VI. 3.3a). It was developed at the Department of Physics and Geology (University of Perugia) and is presented in details in Morgavi et al. (2015) (Figure VI.3.3a; Supplementary Note). The COMMA allows to trigger chaotic mixing processes by controlling the mixing kinematics using the Journal Bearing System (JBS) protocol (Morgavi et al., 2015;[START_REF] Swanson | A comparative computational and experimental study of chaotic mixing of viscous fluids[END_REF]. The setup consists of a Pt80Rh20 crucible (outer cylinder) (77 mm high, 37 mm wide; 𝑅 𝑜𝑢𝑡 = 18.5 mm) filled with the two end-member compositions and an inner spindle located off center (inner cylinder, 𝑅 𝑠𝑝 = 6.2 mm) (Figure VI. 3.3b). Two cylinders were thus drilled out of the crucible containing the dacitic homogenized glass to have the required geometric configuration for the experiment, leaving one empty cylinder for the spindle and a second empty cylinder for the second glassy end-member (synthetic basalt). Then, the crucible is positioned outside the furnace to prepare the required setup for the experiment (see Supplementary Note for more details).

The furnace is brought to the experimental temperature when the crucible is still outside of the furnace (1200 °C) and then moved quickly in the end position required to do the experiment to minimize the time of attainment of thermal equilibrium (~10-15 min) in the sample. The experimental mixing protocol is turned on when the thermocouples inside the spindle measure the experimental temperature. The computer-controlled mixing protocol consists of a combination of separated alternating rotations of outer and inner cylinders in this sequence: two complete rotations of the outer cylinder (crucible) in 7 h and 10 min and six complete rotations of the inner cylinder (spindle) in 3h and 20 min. The rotations of the two cylinders are in opposite directions, which permits to trigger chaotic mixing (streamlines and velocity are time dependent) [START_REF] Swanson | A comparative computational and experimental study of chaotic mixing of viscous fluids[END_REF]. The chaotic magma mixing experiment lasts 10 and a half hours. Mixing and crystallization occurred together at the temperature of 1200 °C, under ambient pressure (1 atm) (Table VI.3.2). The proportion of felsic end-member was higher than the mafic one (80 vol % vs 20 vol %) (Table VI.3.2). At the end of the experiment, rapid quenching is performed to form a glassy experimental product (cooling rate of 41 °C/min) [START_REF] Rossi | Exponential decay of concentration variance during magma mixing: Robustness of a volcanic chronometer and implications for the homogenization of chemical heterogeneities in magmatic systems[END_REF]. After the experiment, the experimental sample was cored out (diameter of 33 mm), and set in epoxy to obtain slices of 2-4 mm thickness. Subsequentially, the patterns generated during the mixing process and the chemical variability between the melts are investigated with a Scanning Electron Microscope and Electron Microprobe Analyzer. experiment) and an estimation of the magma residence time for the natural samples, based on the crystal sizes and experimentally determined plagioclase growth rates. CSD is the population crystal density (defined as the number of crystals in each size interval divided by the size of the interval) versus crystal size [START_REF] Armienti | Decryption of Igneous Rock Textures: Crystal Size Distribution Tools[END_REF][START_REF] Blundy | Petrologic Reconstruction of Magmatic System Variables and Processes[END_REF][START_REF] Cashman | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF][START_REF] Higgins | Measurement of crystal size distributions[END_REF][START_REF] Marsh | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF]. Based on SEM images and crystals sizes, as well as the two-dimensional textural data to threedimensional data, the average crystal aspect ratio of each sample was obtained using a spreadsheet from [START_REF] Morgan | On estimating crystal shape for crystal size distribution analysis[END_REF], CSDSlice (version 5) and the software CSDCorrections (http://www.uqac.ca/mhiggins/csdcorrections.html; version 1.55) was used to obtain information on crystal population density and crystal sizes [START_REF] Higgins | Closure in crystal size distributions (CSD), verification of CSD calculations, and the significance of CSD fans[END_REF][START_REF] Higgins | Measurement of crystal size distributions[END_REF]. Approximately the same number of crystals were studied for the three types of samples (the experimental one, as well as crystals in the andesitic band of the natural banded lavas or the dacitic bands; ~500 crystals) (Supplementary Table VI.3.

1).

𝐺 the growth rate and 𝜏 the residence time were then obtained for the experimental or natural samples, respectively, from the slope of the number density of crystals n depending on the crystal size L, as: 𝑠𝑙𝑜𝑝𝑒 = -1 𝐺 𝜏 (VI. 3.1) (see more details on the method in Supplementary Note).

VI.3.3. Results and discussion: magma mixing textures and timescales in the natural and experimental sample

The experimental samples and the natural samples of Kizimen are not directly compared, as the experiment is mimicking the first mixing process that created the andesite ~50 years ago and the natural banded lavas microlites correspond to the last mixing event with mingling textures.

VI.3.3.1. Textures of Kizimen natural products

In the natural samples, magma mixing evidences are visible on several scales, from macroscopic to microscopic (Figure VI. 3.2), as also evidenced by Auer et al. (2018). At outcrop level, the banded lavas show a complex intermingling between the andesite and dacite, including light grey/dark grey banding features and are either dominated by the dacite or by hybrid andesitic zones and filaments (Figure

VI.3.2a-b).

The compositional difference of the different bands is small (whole rocks of 62 wt% SiO2 dark silicarich andesite and 65 wt% SiO2 light dacitic end-member; Figure VI.3.4), even if a complete separation between the bands was difficult due to the banding/mingling textures (Auer et al., 2018). Study of the residual glasses show that they are rhyolitic in compositions and that the dark glasses are slightly less evolved than in the light bands (Auer et al., 2018;Ostorero et al., accepted)

VI.3.3.2. Eruption triggering timescales

The CSD allows us to study the processes of crystal nucleation and growth and their size distribution, to either obtain their timescales of crystal growth or their nucleation and growth rate.

The natural banded lavas show many phenocrysts, including plagioclase, which are zoned, clustered, blocky and large and which result from several complex magmatic processes, such as magma mixing with a more mafic melt and mingling. Thus, only the microlites in the matrix were studied in the natural banded samples [START_REF] Martel | Eruption Dynamics Inferred from Microlite Crystallization Experiments: Application to Plinian and Dome-forming Eruptions of Mt. Pelee (Martinique, Lesser Antilles)[END_REF][START_REF] Mastin | Insights into volcanic conduit flow from an open-source numerical model[END_REF].

The microlite crystal fraction was estimated at 27 vol% for the andesitic band and at 2 vol% for the dacitic band (Supplementary Table VI [START_REF] Bernard | Controls on eruption style at Rabaul, Papua New Guinea -Insights from microlites, porosity and permeability measurements[END_REF][START_REF] Cashman | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF][START_REF] Marsh | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF]. Timescales of crystal growth can also be derived from the CSD and from experimental growth rates G of plagioclase, using the equation (VI. 3.2):

𝜏 = - 1 𝑠𝑙𝑜𝑝𝑒 𝐺 (VI. 3.2)
We used the experimental growth rates from three experimental studies, either performed with natural basaltic andesitic melts [START_REF] Shea | Kinetics of cooling-and decompression-induced crystallization in hydrous mafic-intermediate magmas[END_REF], dacitic or rhyolitic melts [START_REF] Hammer | An experimental study of the kinetics of decompression-induced crystallization in silicic melt[END_REF][START_REF] Mollard | Decompression-induced Crystallization in Hydrated Silica-rich Melts: Empirical Models of Experimental Plagioclase Nucleation and Growth Kinetics[END_REF]. Average timescales of 12 ± 14 min and of 45 ± 51 min in the andesitic band are found using the maximum and minimum growth rates estimated by [START_REF] Shea | Kinetics of cooling-and decompression-induced crystallization in hydrous mafic-intermediate magmas[END_REF] Comparing these timescales with crystal growth from the previous crystallization experiments, they are shorter than the plagioclase microlites (25-250 µm) growth timescales from [START_REF] Shea | Kinetics of cooling-and decompression-induced crystallization in hydrous mafic-intermediate magmas[END_REF] that were formed over 12 to 48 h [START_REF] Blundy | Ascent-driven crystallisation of dacite magmas at Mount St Helens, 1980-1986[END_REF]. The timescales found in the natural banded samples seem close to skeletal microlites growth of plagioclase (< 15 µm) that formed during decompression-induced crystallization in a few minutes to ~3 h during magma ascent (Arzilli et al., 2019b). The timescales estimated in the banded lavas are also of the same order of magnitude as the ones estimated for mechanical mixing processes in natural magmatic reservoirs, which are of the order of hours [START_REF] Fourmentraux | Crystal fractionation, magma step ascent, and syn-eruptive mingling: the Averno 2 eruption (Phlegraean Fields, Italy)[END_REF]Montagna et al., 2015;Perugini et al., 2010). These timescales of formation of microlites could then correspond to the last magma mixing event and then to the ascent phase before the 2010 eruption of Kizimen. Kizimen is thus an important case study to shed new light on timescales of eruption triggering after magma mixing in a crustal reservoir.

VI.3.3.3. Quantify the kinetics of the first mixing event and crystallization with the experiment

The chaotic high-temperature magma mixing experiment between the synthetic basalt and the natural dacite were also performed to shed light on the first magma mixing process that probably occurred in 1963 at Kizimen, that produced the andesitic magma. VI.3.3.3.1. Textures The starting compositions of the end-members glasses are a basalt and a dacite (Table VI. 3.1;Figure VI.3.4). In the experimental sample, different areas can be identified: i) the dark area, corresponding to the dacite before the experiment, which hereafter will be called "felsic" (Figure VI. 3.7a), ii) the light area, which corresponds to the basalt before the experiment, which will be called "hybrid". Crystallized areas are also present in the basalt and will be called "crystal enclaves", where some felsic enclaves are also sometimes recognized (darker areas in the crystallized enclaves; Figure VI.3.5a-e). The synthetic basaltic end-member was dispersed within the dacitic one by stretching and folding dynamics triggered by the chaotic mixing protocol (Morgavi et al., 2013a) Those present around the crystal enclaves suggest a local enclave disaggregation, as observed in previous experimental studies [START_REF] Laumonier | Experimental simulation of magma mixing at high pressure[END_REF]. Crystal transfer in natural samples is also evidenced and are likely to be accompanied by entrainment of its original melt [START_REF] Cantagrel | Magma mixing: origin of intermediate rocks and "enclaves" from volcanism to plutonism[END_REF][START_REF] Coombs | Magma storage and mixing conditions for the 1953-1974 eruptions of Southwest Trident volcano, Katmai National Park, Alaska[END_REF][START_REF] Gourgaud | Magma mixing in alkaline series: an example from Sancy volcano (Mont-Dore, Massif Central, France)[END_REF][START_REF] Jarvis | Crystal and Volatile Controls on the Mixing and Mingling of Magmas[END_REF]Perugini and Poli, 2012;[START_REF] Ubide | The Relevance of Crystal Transfer to Magma Mixing: a Case Study in Composite Dykes from the Central Pyrenees[END_REF][START_REF] Wright | Heterogeneous pumice populations in the 2.08-Ma Cerro Galán Ignimbrite: implications for magma recharge and ascent preceding a large-volume silicic eruption[END_REF]. The hybrid filaments are all connected to the main mafic area by a thin film of melt (generally between 10-30 µm). The filaments with crystals have a width of ~70 µm and have irregular shapes, showing some deformation with stretched and folded filaments. The interface between the filament and the dacitic host is either sharp or progressive (examples of the progressive margins in Figure VI.3.7b and on the contrary, sharp interfaces in Figure VI.3.7d). The crystals that formed in the hybrid part will be investigated to have more insights into their kinetics of crystallization during the experiment. VI.3.3.3.2. CSD and growth rates The plagioclase crystal fraction was estimated in the experimental sample with the CSD analysis, at 23 vol%, with a tabular dominant shape and mean orientation of -58° (Figure VI. 3.5 3.3, where it indicated abundant growth and limited nucleation, the two types of information are not coherent; this shows the importance of confronting several types of data. CSD results, where 2D data were converted in 3D is more trustworthy so we conclude that nucleation and growth were continuous during crystallization. Extracting the crystal growth rate G from equation 2, it can be assessed by using the slope of the CSD curve and τ, which is the duration of the experiment (10.5 h), if we suppose that the crystals grew from the beginning of the experiment until the end, as well as the nucleation rate (Supplementary Note) [START_REF] Cashman | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF][START_REF] Marsh | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF] [START_REF] Marsh | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF]. The growth rate for the experimental sample is of ~1.10 -2 mm/h and the nucleation rate of 0.25 mm -3 /h (Table VI. 3.3). By using the longest dimension of the crystals and τ, the maximum growth rate, 𝐺 𝐿 is expressed as follows [START_REF] Agostini | Kinetics of crystal evolution as a probe to magmatism at Stromboli (Aeolian Archipelago, Italy)[END_REF][START_REF] Couch | The Kinetics of Degassing-Induced Crystallization at Soufriere Hills Volcano[END_REF][START_REF] Fenn | Nucleation and Growth of Alkali Feldspars From Hydrous Melts[END_REF][START_REF] Hammer | An experimental study of the kinetics of decompression-induced crystallization in silicic melt[END_REF]: 3.3) Using the equation 3, we find a slightly higher growth rate than with the previous equation, of 2.6.10 -2 mm/h and the nucleation rate is of 0.11 mm -3 /h using 𝐺 𝐿 (Supplementary Note).

𝐺 𝐿 = ( 0.5 𝐿 𝑡 𝑒𝑥𝑝 ) (VI.
The growth rate estimates found for the crystals of the experimental sample are of the same order of magnitude (1-2.10 -2 mm/h) as growth rates of plagioclase from decompression experiments in a basaltic andesitic melt (between 1.32.10 -2 mm/h and 1.50.10 -3 mm/h) [START_REF] Shea | Kinetics of cooling-and decompression-induced crystallization in hydrous mafic-intermediate magmas[END_REF], even if no decompression was applied in our experiment (Table VI.3.3). They are slightly higher than those found for experiments with a dacitic (natural starting materials) or rhyolitic melt (synthetic starting materials) (1.02.10 -3 -3.50.10 -6 mm/h and 4. 75.10 -3 -1.75.10 -3 mm/h, respectively) [START_REF] Hammer | An experimental study of the kinetics of decompression-induced crystallization in silicic melt[END_REF][START_REF] Mollard | Decompression-induced Crystallization in Hydrated Silica-rich Melts: Empirical Models of Experimental Plagioclase Nucleation and Growth Kinetics[END_REF]).

Crystal growth rates are variable in natural and experimental systems, as they are dependent on cooling, as well as decompression rates and composition of the melts, they range from 3.6.10 -3 to 3.6.10 -9 mm/h [START_REF] Armienti | Decryption of Igneous Rock Textures: Crystal Size Distribution Tools[END_REF][START_REF] Brugger | Crystal size distribution analysis of plagioclase in experimentally decompressed hydrous rhyodacite magma[END_REF][START_REF] Cashman | Relationship between plagioclase crystallization and cooling rate in basaltic melts[END_REF][START_REF] Cashman | Groundmass crystallization of Mount St. Helens dacite, 1980?1986: a tool for interpreting shallow magmatic processes[END_REF][START_REF] Cooper | Rapid remobilization of magmatic crystals kept in cold storage[END_REF][START_REF] Hammer | Experimental Studies of the Kinetics and Energetics of Magma Crystallization[END_REF][START_REF] Larsen | Experimental study of plagioclase rim growth around anorthite seed crystals in rhyodacitic melt[END_REF][START_REF] Pupier | Experimental quantification of plagioclase crystal size distribution during cooling of a basaltic liquid[END_REF]). The growth rates estimated here are then one order of magnitude less than the maximum of these last rates, so they testify of relatively fast growth during the experiment.

The nucleation rates are between 0.1 to 0.25 mm -3 /h, depending on the growth rate used (Table VI.3.

3).

They are considerably lower than the maximum of the nucleation rates calculated from the experiments discussed before but higher than their minimum nucleation rate (apart from [START_REF] Hammer | An experimental study of the kinetics of decompression-induced crystallization in silicic melt[END_REF]) (3.6.10 -2 -3.6.10 2 mm -3 /s; 90-14.10 3 mm -3 /s; 3.6.10 -2 -42 mm -3 /s, respectively) [START_REF] Hammer | An experimental study of the kinetics of decompression-induced crystallization in silicic melt[END_REF][START_REF] Mollard | Decompression-induced Crystallization in Hydrated Silica-rich Melts: Empirical Models of Experimental Plagioclase Nucleation and Growth Kinetics[END_REF][START_REF] Shea | Kinetics of cooling-and decompression-induced crystallization in hydrous mafic-intermediate magmas[END_REF]. [START_REF] Shea | Kinetics of cooling-and decompression-induced crystallization in hydrous mafic-intermediate magmas[END_REF] evidenced that studies that performed experiments using synthetic starting materials displayed lower nucleation rates and higher growth rates than their equivalent using natural starting melts, which seems to be the case here, even if only one end-member was synthetic.

Our novel mixing experiment at sub-liquidus conditions thus allowed us to reproduce the natural textures of the first mixing process and quantify the kinetics of mixing and crystallization.

Compared to the experiment performed in this study, that lasted 10.5 h, where crystals had time to grow under dynamic conditions and not under pressure, the natural banded andesites seem to have experienced a shorter magmatic event. The triggering process is discussed below.

VI. 3.3.4. Geochemical compositions of the glasses and crystals VI.3.3.4.1. Glasses The composition of the synthetic basalt was almost completely transformed in a hybrid, as it can be seen on Figure VI.3.4: there is a gap in composition between 51 to 55 wt% in SiO2.

The compositions of the filaments as well as those in the crystal enclaves (in the rims, intermediate or cores of these zones; ). This reflects the effects of cross "contamination" between the two compositions, which is driven by crystallization and chemical diffusion [START_REF] Laumonier | Experimental simulation of magma mixing at high pressure[END_REF]Morgavi et al., 2013bMorgavi et al., , 2013a;;[START_REF] Watson | Basalt contamination by continental crust: Some experiments and models[END_REF]. In the residual glasses in the center of these crystallized enclaves, the compositions are more mafic (Supplementary Figure VI.3.5). These areas are less affected by the mixing, or not affected at all, leading to crystallization.

Hybridization then occurred in the experimental sample, by diffusion of elements that led to the composition of a hybrid melt, that replaced the synthetic basaltic one. Thus, as shown before by Morgavi et al. (2013a), chemical diffusion occurred between the filaments, along with stretching and folding processes of mechanical mixing. As chemical elements do not have the same mobility i. VI. 3.3.4.2. Crystals A certain amount of undercooling needs to exist for the crystal nucleation and growth to be initiated (difference between the liquidus Tl and the final T of the system, resting T or Tr) (Mollo and Hammer, 2017). For the experiment here, the degree of undercooling is of 27 ± 0.5 °C (liquidus temperature of the feldspar from MELTS calculations of 1227 °C and final temperature of the system of 1200 ± 0.5 °C). For small undercoolings and a high Tr, interface-controlled growth mechanisms are initiated where diffusion of chemical elements are favored over crystal growth (Mollo and Hammer, 2017): minerals develop euhedral textures (as tabular), as seen in the experiment samples where disequilibrium textures are not present (dendritic, skeletal, sieve-textured, chemical zonations). In such case, the melt composition is homogeneous. The experiment generated a large range of plagioclase compositions, with overlapping An content with natural samples (Supplementary Table VI.3.3). The plagioclase composition predicted by MELTS at 1200°C (An72) is in the main range of anorthite content (𝐴𝑛# = 𝐶𝑎 𝐶𝑎+𝑁𝑎+𝐾 ) of the compositions of the experimental plagioclase (An67-75), meaning that the experimental conditions allowed for the crystallization of plagioclase of this composition on a wider range, as the compositions describe An52- The experimental plagioclase crystals show a slight normal zonation (Na higher on the rims, Supplementary Figure VI.3.7), with the cores of the crystals that are mainly between An60-76 whereas the rims are lower in An#, mainly between An50-63. This normal zoning could come from the quenching of the experiment, with the abrupt drop in temperature at the end of the experiment, as temperature changes influence the composition of plagioclase. Indeed, the An content of plagioclase decreases with decreasing temperature, developing normal zoning [START_REF] Ustunisik | New insights into the processes controlling compositional zoning in plagioclase[END_REF]. Waters & Lange (2017) also proposed that progressive loss of H2O (due to degassing of the magma upon ascent) can explain the presence of normally zoned plagioclase crystals in a crystal-poor silicic liquid, but this hypothesis was not tested during the experiment as starting materials were anhydrous.

VI.3.3.5. Magma dynamics from the experiment and natural samples

The experiment allowed us, for the first time, to study crystallization and mixing, at fixed temperature and mimics the first mixing process that could have occurred in 1963. The timing of the experiment also gives us an idea of the time it takes to homogenize the basaltic end-member, taking less than 10.5 h but leaving some evidence of the mixing process (Figure VI. 3.11a). The basaltic end-member can thus not be sampled in the deposits, due to fast mixing. Growth rates from the experiment also shows that the crystallization process was fast, as it was high compared to experimentally growth rates for plagioclase. This mixing event could then have occurred sometime before the 2010 eruption of Kizimen, without triggering the eruption ( One scenario to explain the triggering of the eruption would be a third and last magma mixing phase, due to the magma mingling evidence summarized here: inhomogeneous banded lavas, compositional bands with different whole rock compositions and slightly different residual glasses, different textures of the matrix of the different bands (the andesitic one being richer in microphenocrysts and microlites) and disequilibrium textures. This magma mixing phase could have then generated some decompression, destabilizing the magmas, leading to the development of slight normal zoning in the plagioclase and to the eruption (Figure VI. 3.11c).

Using the timescales studied here with the CSD on the natural samples, we can decipher the last magmatic event that occurred before the eruption. According to the CSD in the natural samples, timescales of the order of minutes to hours or even days are found for the timescales of residence of the microlites in the magmas, when considering the same experimental growth rates for plagioclase (Table VI.3.3), with two phases of microlites formation. As both the microlites in the natural dacitic and andesitic bands record first a nucleation event, followed by a growth-dominated regime, the nucleation event could correspond to the onset of the last magma mixing phase and then the decompression event occurred, generating the microlites of the growth phase, in a few minutes (Figure VI.3.11c). This kind of event is recognized as a triggering event before eruptions (Arzilli et al., 2019b;[START_REF] Ginibre | Crystal Zoning as an Archive for Magma Evolution[END_REF].

VI.3.4. Conclusion

To conclude, in this study, a chemical mixing at sub-liquidus conditions via a chaotic mixing experiment was performed for the first time to better understand the mixing event at the origin of the magmas before the 2010 eruption of Kizimen. This experiment was done to reproduce the complex natural case and understand how the magma mixing processes influenced the loss of the basaltic endmember in the erupted products (Figure VI.3.11). The experiments show chaotic textures of magma mixing, with a homogenization of the basaltic end-member in a few hours. The triggering process before the 2010 eruption of Kizimen was investigated by looking at natural banded lavas. The last magma mixing/mingling event occurred before the eruption, in the order of minutes to hours, followed by a decompression event in minutes leading to the eruption (Figure VI. 3.11). Magma mixing events are common events prior to many eruptions of many arc volcanoes (such as at Mt Unzen (Japan), Pinatubo (Philippines), Mt Dutton (Alaska), Lassen Peak (California) or Soufrière Hills (Montserrat)). Some key questions still remain, such as understanding which factors determine if a magmatic injection leads to an eruption or also what is the role of crystals during magma mixing processes. Some simplifications of the natural system were necessary here, but we were able to understand that crystallization and mixing occurred together during the experiment and influenced each other. It is thus fundamental to have a better understanding of the roles of crystals and mixing acting together before volcanic eruptions. New experiments focused on studying the effects of crystals on mixing would be interesting to perform.

(i) 2 complete rotations of the outer cylinder in 7 h and 10 min (ii) 6 complete rotations of the inner cylinder in 3h and 20 min.

The mixing experiment lasted ten and half hours. The main fluid-dynamical parameters characterizing COMMA experiments are summarized in Table VI. 3.2. At the end of the experiment, the furnace holding the experimental crucible was quickly moved vertically while maintaining the crucible and the spindle in their original positions (Morgavi et al., 2015;[START_REF] Rossi | Exponential decay of concentration variance during magma mixing: Robustness of a volcanic chronometer and implications for the homogenization of chemical heterogeneities in magmatic systems[END_REF]. This final step allows a rapid quenching of the experimental melt (cooling rate of 41 °C/min) to form a glassy experimental product [START_REF] Rossi | Exponential decay of concentration variance during magma mixing: Robustness of a volcanic chronometer and implications for the homogenization of chemical heterogeneities in magmatic systems[END_REF]. After the experiment, quenched glasses are cored out from the outer cylinder and set in epoxy. They are then cut into slices of 2-4 mm thickness and analyzed by Scanning Electron Microscope and Electron Microprobe Analyzer to study the patterns generated by the mixing process and the space and time variability of chemical exchanges between the melts. VI.3.A.1.4. Crystal Size Distribution SEM images were analysed with ImageJ software (https://imagej.nih.gov/ij/; version 1.52a), applying a threshold on the greyscale to select the microlites and then displaying the analysed particles as a mask.

Images were then processed with JMicrovision (https://jmicrovision.github.io/; version 1.3.4) to measure the length, width, area of microlites and orientation (vesicle-free), with the edge-intersecting crystals that were removed. To convert two-dimensional textural data to three-dimensional data, the average crystal aspect ratio of each sample was obtained using a spreadsheet from [START_REF] Morgan | On estimating crystal shape for crystal size distribution analysis[END_REF], CSDSlice (version 5). This aims to compare the 2D measurements with 2D ideal shape curves from a database of thousands of random sections through known shapes, to identify five best-fitting curves and the corresponding crystal aspect ratios based on a least-square fit between the sample and the database [START_REF] Bernard | Controls on eruption style at Rabaul, Papua New Guinea -Insights from microlites, porosity and permeability measurements[END_REF][START_REF] Morgan | On estimating crystal shape for crystal size distribution analysis[END_REF]. More than 400 crystals were measured for each sample [START_REF] Morgan | On estimating crystal shape for crystal size distribution analysis[END_REF]. Out of the three samples, the R 2 of the best-fit was over 0.8, meaning a reliable statistical fit (Supplementary Table VI.1).

The crystal fraction was calculated on a vesicle-free basis: 𝜑 = 𝐴 𝑐𝑟𝑦𝑠𝑡𝑎𝑙 𝐴 𝑇 (IV. 3. A. 1) Where 𝐴 𝑐𝑟𝑦𝑠𝑡𝑎𝑙 is the area of the crystal phase and 𝐴 𝑇 is the total area of the images studied. The twodimensional crystal number density 𝑁 𝑎 (mm -2 ) was obtained by dividing the number of microlites by the total area analysed, free of vesicles [START_REF] Blundy | Petrologic Reconstruction of Magmatic System Variables and Processes[END_REF].

The software CSDCorrections (http://www.uqac.ca/mhiggins/csdcorrections.html; version 1.55) was then used to obtain information on crystal population density and crystal sizes [START_REF] Higgins | Closure in crystal size distributions (CSD), verification of CSD calculations, and the significance of CSD fans[END_REF][START_REF] Higgins | Measurement of crystal size distributions[END_REF].

The number density of crystals n is a linear function of crystal size L, expressed as follows:

𝑛 = 𝑛 0 𝑒 -𝐿 𝐺 𝜏 (IV. 3. A. 2)
Where 𝑛 0 is the nucleation density of the curve of the crystal size distribution, for L = 0, 𝐺 the growth rate and 𝜏 the residence time [START_REF] Higgins | Measurement of crystal size distributions[END_REF]. This curve, when plotted in logarithmic, gives us access to the timescales of magmatic processes by using the linear relation between the crystal population density and the crystal length, as the slope of the correlation is equal to: et al., 2019b;[START_REF] Higgins | Measurement of crystal size distributions[END_REF].

𝑠𝑙𝑜𝑝𝑒 = - 1 𝐺 𝜏 (IV. 3. A. 3) (Arzilli
The nucleation rate can also be calculated, using the intercept of the CSD curve or nucleation density, n 0 , which is the ratio of nucleation rate over growth rate (J/G) [START_REF] Cashman | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF][START_REF] Marsh | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF].

The three-dimensional crystal number density 𝑁 𝑣 (mm -3 ) was obtained with CSDCorrections, which adds the number of crystals in each bin interval after the stereological correction with CSDSlice. The characteristic length of the microlites LC, is also estimated by CSDCorrections (𝐿 𝐶 = -

VI.3.A.2.3. Near crystallized areas

The glasses in the crystallized enclaves also show variations, with the cores of these zones being more mafic than the margins (mainly between 53-60 wt% compared to 56-64 wt% in SiO2) (Supplementary Figure ) of the crystals have been investigated as well as their textures (Supplementary Table VI.3.3 The predicted An# by MELTS software for the synthetic basalt at 1200°C and NNO + II is shown for comparison (An73) [START_REF] Gualda | Rhyolite-MELTS: a Modified Calibration of MELTS Optimized for Silica-rich, Fluid-bearing Magmatic Systems[END_REF]. The cores of the crystals are then close to the predicted An#. Some clinopyroxenes are also found (analyses in several crystals, Supplementary Figure VI.3.6) The aims of the second experiment carried out were to study the natural system of Kizimen (as the endmember compositions were close from the mixing process identified shortly before the eruption (mixing between the andesite and dacite)) and the crystallization process during mixing.

VI.4.1. Material and methods

The two-natural end-members were the dacite from the 2010-2013 eruption and a basaltic andesite (Kizi-96-24, dating from 1200-1300 years ago) (Churikova et al., 2013b(Churikova et al., , 2007)). The objectives were to produce 10-15 vol% of feldspar in the two different melts at ambient pressure before mixing and then to mix these two starting materials during the experiment, allowing crystallization of the basaltic andesite (at a temperature under the liquidus of feldspar for the basaltic andesite and above for the dacite). Following MELTS simulations, the basalt evolves in a basaltic andesite, as emitted by Kizimen, at 1172 °C and 13 vol% of feldspar (Table VI.4.1) [START_REF] Gualda | Rhyolite-MELTS: a Modified Calibration of MELTS Optimized for Silica-rich, Fluid-bearing Magmatic Systems[END_REF]. Dacitic and basaltic andesitic powders were melted and homogenized for 1 h at 1500 °C and then during 6h at 1115 or 1172 °C inside a high-temperature furnace (Nabertherm® HT 04/17) at ambient pressure at the Department of Physics and Geology (University of Perugia) in Pt80Rh20 crucibles (Morgavi et al., 2015).

Based on MELTS calculations of the crystal fraction for both end-members (Table VI. 4.1), we decided to perform the experiment at a temperature below the liquidus of the feldspar and spinel for the basaltic andesite (around 1203 and 1223°C respectively, at NNO + II conditions) and above them for the dacite (around 1150 and 1153°C respectively, at NNO + II conditions). The experiment was then run at a temperature of 1180 °C, at ambient pressure for 10.5 h. At 1200°C, a crystal fraction of 14 vol% in the basaltic andesite is expected (with 10 vol% of feldspar) [START_REF] Gualda | Rhyolite-MELTS: a Modified Calibration of MELTS Optimized for Silica-rich, Fluid-bearing Magmatic Systems[END_REF].

T (°C)

Vol % of crystals fd (Melts)

VI.4.4.2. Strain influence on crystallization

Time evolution studies of magma mixing processes show that fractal figures are generated, due to stretching and folding of magmas [START_REF] De Rosa | Fractal analysis of mingled/mixed magmas: an example from the Upper Pollara eruption (Salina Island, southern Tyrrhenian Sea, Italy)[END_REF][START_REF] Perugini | Chaotic advection, fractals and diffusion during mixing of magmas: evidence from lava flows[END_REF][START_REF] Wada | Fractal structure of heterogeneous ejecta from the Me-akan volcano, eastern Hokkaido, Japan: implications for mixing mechanism in a volcanic conduit[END_REF]. As in our experiment, the first rotation generated some stretching, it led to areas where crystals grew and then, the second rotation generated convection with the involved crystals that are dispersed in the dacite. Indeed, isothermal deformation experiments or cooling deformation experiments show that an increasing deformation rate promotes nucleation and crystal growth by decreasing the nucleation time (increase of the random collisions between atoms) and by fresh melt supply to the growing crystals [START_REF] Di Fiore | An Extended Rheological Map of Pāhoehoe-'A'ā Transition[END_REF][START_REF] Kolzenburg | The effect of oxygen fugacity on the rheological evolution of crystallizing basaltic melts[END_REF][START_REF] Kolzenburg | In situ thermal characterization of cooling/crystallizing lavas during rheology measurements and implications for lava flow emplacement[END_REF][START_REF] Kouchi | Effect of stirring on crystallization kinetics of basalt: texture and element partitioning[END_REF][START_REF] Vona | The effects of undercooling and deformation rates on the crystallization kinetics of Stromboli and Etna basalts[END_REF].

A recent study show that high initial degrees of undercooling and shear rates enhance crystallization by reducing the incubation time of nucleation [START_REF] Di Fiore | An Extended Rheological Map of Pāhoehoe-'A'ā Transition[END_REF]. Shear rates from 1-20 s -1 were tested while the shear rate expected to occur at depth are several orders of magnitude lower, around 10 -7 s -1 [START_REF] Williams | Microgranitic enclave shapes and magmatic strain histories: Constraints from drop deformation theory[END_REF] or from 10 -12 to 10 -2 s -1 [START_REF] Albertz | Fast strain rates during pluton emplacement: Magmatically folded leucocratic dikes in aureoles of the Mount Stuart Batholith, Washington, and the Tuolumne Intrusive Suite, California[END_REF][START_REF] Di Fiore | An Extended Rheological Map of Pāhoehoe-'A'ā Transition[END_REF][START_REF] Laumonier | Experimental simulation of magma mixing at high pressure[END_REF]. Kouchi andSunagawa (1982, 1985) experiments were performed applying 48 rotations per minute (during 2 h), whilst De Campos et al. ( 2011) applied a minimum rate of about 4 tours every hour which the latter being broadly equivalent to a strain rate of 10 -1 s -1 [START_REF] Laumonier | Experimental simulation of magma mixing at high pressure[END_REF]. For our experiment, the mixing protocol consisted in separated alternating rotations of the outer and inner cylinders, leading to separate strain rates.

VI.4.5. Perspectives

To conclude, this experiment could allow to focus on the effects of crystals during mixing and to study the fractal patterns generated during the experiment. Some additional work would consist in the more detailed study of the chemical variations generated to know more about the amount of differentiation. The fractal nature of the mixing patterns could be shown by using the box technique method (relationship between a square mesh of size r laid over the image and the number of boxes Nr containing the pixels belonging to the interface between the two melts) [START_REF] Mandelbrot | The Fractal Geometry of Nature[END_REF]Morgavi et al., 2013a).

Analyses by ICP-MS Laser Ablation to determine the trace element variation in the experimental samples are planned to study the extent of diffusive fractionation of trace elements, which depends on the mixing time. This approach could also be applied to the natural case study of the banded lavas of Kizimen and allowed us to have an estimate of the mixing time elapsed between the start of the mixing processes and the eruption (Morgavi et al., 2013b;Perugini et al., 2010[START_REF] Perugini | Trace element mobility during magma mixing: Preliminary experimental results[END_REF][START_REF] Perugini | Diffusive fractionation of trace elements by chaotic mixing of magmas[END_REF].

As in Chapter VI.3, Crystal Size Distribution could also be performed in the experimental samples to know the growth kinetics of the crystals [START_REF] Cashman | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF][START_REF] Marsh | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF]. Other experiments could be done by studying magma mixing with various mixing times, shorter or longer than what was done for the two experiments (Chapters VI.3-4) performed to compare the mixing patterns depending on the mixing time.

Résumé

Cette thèse a été réalisée sur des volcans d'arc de subduction : au niveau de l'île de la Dominique (Petites Antilles, principalement sur le volcan Morne Trois Pitons-Micotrin) et au Kamchatka (Russie, sur le Bezymianny et Kizimen). Sept éruptions de ces volcans ont été étudiées, qu'elles soient non monitorées (5 éruptions à la Dominique allant de 9 à 24 ka) ou monitorées (l'éruption bien surveillée de 2010-2013 du Kizimen et l'éruption de 1956 du Bezymianny ; même si une seule station sismique distante était présente à l'époque (à 42 km) et que les observations étaient difficiles en raison des conditions de terrain). Ces éruptions ont été choisies pour i) suivre l'histoire pré-éruptive du stockage du magma dans les réservoirs magmatiques ; ii) contraindre les échelles de temps des processus pétrologiques impliqués ; iii) pour les volcans monitorés, identifier si les échelles de temps de diffusion peuvent être corrélées à l'activité de surveillance enregistrée ; iv) comprendre comment le mélange et le brassage se produisent avant une éruption, par une étude expérimentale.

Les cristaux d'orthopyroxènes ont été les plus utilisés car ils sont sensibles aux conditions de stockage magmatique et peuvent donc être utilisés pour étudier les processus magmatiques. Pour les volcans du Kamchatka, une combinaison de la détermination de ces échelles de temps avec plusieurs minéraux des mêmes échantillons a été menée afin d'obtenir des modèles complets de dynamiques des magmas avant les éruptions étudiées.

Sur l'île de la Dominique, l'étude intégrée sur plusieurs éruptions successives générées au cours des 24 000 dernières années par différentes zones de stockage à différentes profondeurs dans un système transcrustal a mis en évidence des phénomènes de mélanges des magmas plusieurs dizaines d'années avant les éruptions, avec des mélanges plus marqués quelques années avant les éruptions Pliniennes (Chapter IV). Les grandes éruptions ponceuses (ignimbritiques) semblent avoir des échelles de temps plus longues que les éruptions Pliniennes. Ces résultats d'échelles de temps peuvent être utiles pour le suivi des signaux de réactivation enregistrés en surface, afin d'anticiper le temps probable restant avant une éruption dans le cas de la réactivation de Morne Trois Pitons-Micotrin, ce dernier n'ayant pas connu d'éruption magmatique historique (donc monitorée).

Au Kamchatka, le décryptage des dynamiques pré-éruptives est particulièrement intéressant pour comprendre les processus magmatiques précédant les éruptions majeures de cette péninsule, car cet arc est l'un des arcs volcaniques les plus productifs au monde (Churikova et al., 2001;Melekestsev et al., 1995;[START_REF] Ponomareva | Late Pleistocene-Holocene volcanism on the Kamchatka Peninsula, Northwest Pacific Region, in: Volcanism and Subduction: The Kamchatka Region[END_REF]. L'étude du Bezymianny a fourni de nouvelles perspectives temporelles sur les dynamiques pré-éruptives des magmas avant la phase paroxysmale de l'éruption du Bezymianny, ayant généré une explosion latéralement dirigée. Un stockage des magmas dans un cryptodôme a débuté quelques mois avant l'effondrement d'un flanc du volcan conduisant à un blast, suivi de la dépressurisation du magma stocké plus en profondeur dans la croûte et ayant conduit à une phase Plinienne (Chapter V). Pour le Kizimen, une bonne corrélation entre les ensembles de données pétrologiques (orthopyroxènes, magnétites et le dernier événement enregistré par les plagioclases) et les signaux de surveillance géophysique a été obtenue pour les magmas évolués du Kizimen (Chapters VI.1-2). Cette corrélation est une avancée majeure pour la surveillance volcanique dans les zones de subduction et pourrait aider à développer des horloges d'alerte précoce avant les éruptions, car ces types de corrélations sont moins étudiés pour les magmas évolués, par rapport aux systèmes mafiques. L'étude expérimentale menée dans le cadre de projets de mobilité internationale ont permis d'apporter des contraintes temporelles sur un mélange entre deux magmas de compositions différentes sous le Kizimen ainsi que sur la phase de mélange rapide ayant eu lieu quelques minutes ou quelques heures avant l'éruption (Chapters VI.3-4).

Malgré certaines limitations concernant l'utilisation d'un coefficient de diffusion spécifique ou des incertitudes liées aux échelles de temps ou simplifications dans le cadre des expériences de mélanges des magmas, des échelles de temps de processus magmatiques ont été obtenues pour ces éruptions explosives. Ces temps permettent de mieux comprendre les systèmes magmatiques et leur réactivation. De plus, la comparaison de ces temps avec des signaux géophysiques et géochimiques de surveillance des volcans par les observatoires nous amènent à pouvoir les intégrer dans des futurs modèles de scénarios éruptifs.

Pour passer de ces études de diffusion à une horloge d'alerte précoce, il faut étudier plusieurs éruptions du même volcan pour voir si différentes parties du système ont été réactivées entre les éruptions et si elles enregistrent la même histoire au cours des éruptions. Cette étude de la présence ou non d'une systématique de la dynamique pré-éruptive est vraiment nécessaire pour pouvoir appliquer ce modèle de réactivation magmatique à d'autres volcans de zones de subduction qui n'ont pas produit d'éruptions magmatiques historiques.

Les perspectives principales de ce travail seraient de déterminer un nouveau coefficient d'interdiffusion pour les orthopyroxènes, réalisé pour des compositions semblables à celles étudiées ici ou dans les volcans d'arc et en contraignant la fugacité d'oxygène. Cette meilleure contrainte sur la dépendance par rapport à la fugacité d'oxygène permettrait de voir si les temps obtenus seraient effectivement plus courts. L'étude d'autres éruptions passées monitorées de volcans d'arc de subduction serait aussi souhaitable, afin de comparer les dynamiques pré-éruptives et déterminer des temps restants avant une éruption selon la détection de signaux spécifiques en surface. Étudier les processus de remontée des magmas et de taux de décompression serait aussi fondamental afin de coupler les dynamiques pré-éruptives du réservoir avec celles dans les conduits magmatiques, dont le style éruptif d'une éruption dépend grandement.

VII.1. Pre-eruptive magma dynamics and timescales

VII.1.1. Main results of this study

This thesis was carried out on arc volcanoes: in Dominica (Lesser Antilles; mainly working on Morne Trois Pitons-Micotrin) and in Kamchatka (Russia; on Bezymianny and Kizimen). Seven eruptions of these volcanoes were studied, both unmonitored (5 eruptions in Dominica ranging from 9 to 24 ka) or monitored (the well-monitored 2010-2013 eruption of Kizimen and the less monitored 1956 climactic phase of the eruption of Bezymianny; only one remote seismic station was present at the time (42 km away from the volcano) and observations were difficult due to the field conditions). These eruptions were chosen to: i)

Track the pre-eruptive history of magma storage in the magmatic plumbing system ii)

Constrain the timescales of the petrological processes involved iii)

For monitored volcanoes, to identify if the diffusion timescales could be correlated to the monitoring activity recorded iv)

To understand how mixing and/or mingling occur prior to an eruption, by an experimental study Orthopyroxenes crystals (opx) carried by the magmas involved in these 7 eruptions were the most studied. Why? Opx are major rock-forming minerals in the Earth's upper mantle and lower crust and because of their unique crystal structures and major element compositions, they can be used to study magmatic processes. Their compositions are thus sensitive to modifications of the magma composition (T, P, fH2O, fO2…) making them reliable archives of magma dynamics.

Other minerals from the same eruptions, such as magnetites for Bezymianny and Kizimen, as well as plagioclase crystals for Kizimen were also used. The zoned crystals recorded the magmatic processes that the magmas experienced prior to the eruption (provided that the chemical gradients were not homogenized when the eruption occurred). Melt inclusions and residual glasses also gave us some insights in the compositions of the pre-eruptive melts and residual melts involved during the magmatic/volcanic processes.

Pre-eruptive timescales in the order of a few days to several decades have been modelled in these 7 eruptions (Chapters IV-VI).

In Dominica, we performed an integrated study on several successive eruptions generated in the last 24 kyrs by different storage areas at different depths in a transcrustal system (Chapter IV).

For the Morne Trois Pitons-Micotrin volcano, we proposed a model of rejuvenation of the reservoirs by mixing processes mostly in the last decade prior to the eruptions, by coupling a complete textural and compositional analysis of orthopyroxenes using the Crystal System Analysis, an underutilized technique up to now, and diffusion chronometry (2410 opx mounted). This study is confirmed by melt inclusions and matrix glasses investigation from previous studies. We also showed that there was a systematic pattern of the estimated pre-eruptive timescales and pre-eruptive dynamics for Plinian eruptions whose magmas are stored at the same depth (from 2 to 8 km [START_REF]Evidence for an Active, Transcrustal Magma System in the Last 60 ka and Eruptive Degassing Budget (H 2 O, CO 2[END_REF]. Crystal system analysis and timescales of zoned orthopyroxenes show that mixing between two main magmatic environments in the 10 to 30 years prior to the eruption occurred. The interactions between the magmas accelerated in the last 2 years prior to the eruption. The estimated timescales for the large pumiceous (ignimbritic) eruption of Grand Fond, where magmas were stored deeper (from 12 to 16 km (Solaro et al., 2019)), are different from those for Plinian eruptions, with magma mixing beginning a century before the eruption. For other large eruptions of Morne Trois Pitons-Micotrin volcano (Roseau, 33 ka) and of the neighboring Morne Diablotins volcano (Layou, 55 ka), stored deeper (12-16 km (Solaro et al., 2019)), two magma batches partially mixed, also during several decades [START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF] (Appendix A). Plinian eruptions and large ignimbritic eruptions at these volcanoes then seem to have recorded similar magmatic processes with convection and magma mixing processes in a thermally zoned reservoir for the Plinian eruptions and Grand Fond and heating of the magmas by a magma injection (for Layou and Roseau), with timescales of several decades to some years. Large eruptions seem to have longer timescales than Plinian eruptions. These timescales results can be useful for the monitoring of reactivation signals recorded on the surface, in order to anticipate the likely time remaining before an eruption in case of the reactivation of Morne Trois Pitons-Micotrin, as no monitored eruption exists.

In Kamchatka, deciphering the pre-eruptive dynamics is of particular interest to understand the magmatic processes prior to major eruptions of this peninsula, as this arc is one of the world's most productive volcanic arc and cluster of subduction volcanoes (Churikova et al., 2001;Melekestsev et al., 1995;[START_REF] Ponomareva | Late Pleistocene-Holocene volcanism on the Kamchatka Peninsula, Northwest Pacific Region, in: Volcanism and Subduction: The Kamchatka Region[END_REF]. Constrain magma movements and their spatiotemporal dynamics of two volcanoes from the Central Kamchatka Depression was done in this thesis. Bezymianny is part of the Klyuchesvkoy group, not far away from Kizimen, with Kizimen magmas that may have a similar origin to those from this volcanic group. The first historical magmatic eruptions of these volcanoes were studied, which occurred after a repose time of 1000 years for Bezymianny and at least 80 years for Kizimen (Churikova et al., 2007). Bezymianny 1956 climactic phase of the eruption is emblematic as this eruption was among the greatest explosive ones of the 20 th century and because it produced a "laterally directed blast", devastating an area of 500 km 3 around the volcano, generating 1.4-1.7 km 3 of eruptive ejecta throughout several eruptive phases (Table VII.1) [START_REF] Belousov | Deposits of the 30 March 1956 directed blast at Bezymianny volcano, Kamchatka, Russia[END_REF].

We decided to use several minerals in each study of Bezymianny and Kizimen, to have a complementary view on the magmatic processes and their timescales, as their diffusivities are different and as few studies have modelled diffusion in multiple crystals from the same eruption [START_REF] Brugman | Common assumptions and methods yield overestimated diffusive timescales, as exemplified in a Yellowstone post-caldera lava[END_REF]Chamberlain et al., 2014;[START_REF] Cooper | Rapid assembly and rejuvenation of a large silicic magmatic system: Insights from mineral diffusive profiles in the Kidnappers and Rocky Hill deposits, New Zealand[END_REF]Fabbro et al., 2018;Flaherty et al., 2018;[START_REF] Magee | The Lead-up to Mount Etna's Most Destructive Historic Eruption[END_REF][START_REF] Shamloo | Decadal transition from quiescence to supereruption: petrologic investigation of the Lava Creek Tuff, Yellowstone Caldera[END_REF]Singer et al., 2016) and modelled both the inner and outer rims in multiple zoned crystals to understand the complete pre-eruptive dynamics recorded by the crystals, as we did in Dominica (Figure VII.1).

Based on the zoning patterns of orthopyroxenes and magnetites of volcanic clasts from the 1956 climactic phase of the eruption of Bezymianny (blast and Plinian deposits), as well as residual glasses and melt inclusions, we present that Bezymianny system was reactivated, at least 2-3 years prior to the climactic phase, with the majority of the timescales below 6 months (Chapter V). The magma is approximatively located between 9-15 km and convection in this magma reservoir occurred. Then, for the magma at the origin of the laterally directed blast, magma ascent and storage in the cryptodome happened and then some degassing or heating of the magma, that could have come from the injection of a hotter magma at depth. Storage in the cryptodome was correlated with deformation of the southeastern flank of the volcano. This study thus provided new timescales insights into the pre-eruptive dynamics of the magmas prior to the paroxysmal phase of the eruption of Bezymianny, considering all crystals of the crystal assemblage studied in the literature to have a complete model of these magma dynamics.

For Kizimen, two approaches were used: a study of the natural samples via the compositions and zoning patterns of several crystals and experimental studies to have a complete view of the magma dynamics. Considering these two different approaches, the timescales of several phases were inferred:

-In 1963 (~50 years before the eruption), the andesitic magma from Kizimen could have been formed in a few hours following the injection of mafic magma in a dacitic reservoir (based on the chaotic magma mixing experimental study; Chapter VI.3). -1.5 years before the eruption of 2010, magma mixing in a stratified magma chamber (andesitic in the lower part and dacitic in the upper part) was inferred to have occurred, up to the time of eruption. It led to interactions between three main magmatic environments, considering three types of crystals (opx, plag and mgt) in two types of volcanic clasts (Chapters VI.1-2). Dacite recorded an evolution towards a colder magmatic environment while the andesites interacted with a hotter magmatic environment. Other events may have been recorded by plag, which could be related to the seismic activity before 2009. -In minutes to hours before the eruption, the mingling event that potentially triggered the eruption occurred (natural banded lavas study; Chapter VI.3). The second experiment could also allow to estimate the timescales of the mingling phase, as the compositions of the magmas involved are close (Chapter VI.4). These timescales can be a valuable information on the triggering mechanism before the eruption as magma mingling was identified as a triggering process of the eruption and can be used for future monitoring of eruptions, knowing the time between the magmatic event and the eruption. The seismic monitoring data for Kizimen was of better quality than Bezymianny, as Kizimen is isolated from other volcanoes and only the eruption of Kizimen was recorded in the seismic signals. A good correlation between petrological datasets (opx, mgt and the last event recorded by the plag) and geophysical monitoring signals was obtained for the evolved magmas of Kizimen. This correlation is a major advance for volcanic monitoring in subduction zones and could help develop early-warning clocks prior to eruptions as these kinds of correlations are less studied for evolved magmas, compared to mafic systems (Albert et al., 2019;[START_REF] Kahl | Compositionally zoned crystals and real-time degassing data reveal changes in magma transfer dynamics during the 2006 summit eruptive episodes of Mt. Etna[END_REF]Kahl et al., , 2011;;Longpre et al., 2014;[START_REF] Marti | Correlation of Magma Evolution and Geophysical Monitoring during the 2011-2012 El Hierro (Canary Islands) Submarine Eruption[END_REF]Pankhurst et al., 2018;[START_REF] Rae | Time scales of magma transport and mixing at Kīlauea Volcano, Hawai'i[END_REF]Sundermeyer et al., 2020a).

Comparing the magma dynamics between the eruptions at Bezymianny and Kizimen, the magmatic processes seem to have occurred mainly in the last 2 years prior to the eruptions. It can be highlighted that these pre-eruptive dynamics are shorter than the magma mixing processes deciphered for the pumiceous eruptions of Dominica (several years compared to several decades), even if the magma dynamics for Plinian eruptions of Dominica accelerated in the last two years prior eruption (Table VII.1). Between these seven eruptions, the estimated VEI were of 3 for Kizimen 2010 eruption, 4 for the Plinian eruptions of Dominica, and 5 for Grand Fond ignimbritic eruption in Dominica and Bezymianny 1956 climactic phase [START_REF] Newhall | Mechanisms of igneous intrusion[END_REF], with different magma volumes; Grand Fond erupted magma volumes being the more voluminous (Table VII.1). These eruptions are comparable in terms of magma compositions (andesite/dacite) while the magma storage depths are different, with magmas at the origin of Plinian eruptions of Dominica that are stored shallower (Table VII.1). Magmas giving Grand Fond eruption and Bezymianny 1956 one have comparable storage depths (12-16 km compared to 9-15 km; Table VII.1). The magmatic reservoirs were reactivated, either by magmatic injections that participated in heating the reservoir (Grand Fond and Bezymianny 1956), or magma mixing in a temperature-zoned reservoir (Goodwill Entrance, PPR1-3) or between different magma compositions (andesite/dacite for Kizimen 2010). The En content of opx investigated was close with more than 40 opx that have been investigated for each eruption and between 9 and 69 timescales that have been modelled for each eruption (Table VII.

1).

All these timescales obtained on the different eruptions could be used in case of a future eruption on some of these volcanoes, in the case of the detection of similar unrest signals and could improve early warning detection, eruption forecasting and management of volcanic crisis for different eruptive scenarios. From the correlation obtained for Kizimen, we could think that characteristic unrest would be a seismic crisis with earthquakes located from the base of the magmatic reservoirs to the surface (magnitudes < 4) [START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF] with some ground deformation that could also be measured. Early-warning clocks could be the next step of this type of studies, for a specific volcano. In case of volcanic unrest on Morne Trois-Pitons volcano or on Kizimen, comparison between different types of signals can be undertaken, such as deformation, seismic, or temperature signals to compare them with petrological timescales deciphered here. For Bezymianny, the volcano is still very active but the dynamics changed since 1956, leading to more open-system eruptive dynamics and one or two eruptions occurring per year. The good correlation of the timescales with monitoring signals at Kizimen is promising for this system and the application of these types of correlations will be used for unmonitored eruptions in other volcanic arcs, such as the Lesser Antilles volcanoes, which was one of the objectives of the ANR project in which my PhD thesis is integrated. The pre-eruptive timescales can be precisely compared when the volcanic context is the same (for example, hot spots or volcanic arcs) and for volcanic arcs, when the magma accumulation phenomena are approximately the same. However, a general comparison between timescales of eruptions involving the same magma compositions can be done. The timescales of volcanoes from mid-ocean ridges and hot spots involving basaltic (basanite or tephrite magma compositions) are relatively short: they are all below half a year (the median of the longest timescale of this type of system is of ~130 days, for the East Pacific Rise) [START_REF] Pan | Mid-ocean ridge magma chamber processes: Constraints from olivine zonation in lavas from the East Pacific Rise at 9°30′N and 10°30′N[END_REF] and for the volcanoes from hot spots or Icelandic volcanoes (context of both hot spot and mid-ocean ridge), around 1 year (median of ~400 days days for Kīlauea 1959 and Bárðarbunga Early Holocene) [START_REF] Caracciolo | Timescales of crystal mush mobilization in the Bárðarbunga-Veiðivötn volcanic system based on olivine diffusion chronometry[END_REF]Couperthwaite et al., 2020)) (Figure VII.1). Bezymianny and Kizimen median timescales are plotting in the same timescales ranges as these less evolved systems (mostly below 1 year), even if other volcanoes from subduction zones with evolved compositions also give short timescales, such as Auckland Volcanic Field, Ruapehu (New Zealand; boundary between Pacific and Australian plates) [START_REF] Brenna | Olivine xenocryst diffusion reveals rapid monogenetic basaltic magma ascent following complex storage at Pupuke Maar, Auckland Volcanic Field, New Zealand[END_REF][START_REF] Kilgour | Timescales of magmatic processes at Ruapehu volcano from diffusion chronometry and their comparison to monitoring data[END_REF], Novarupta (Katmai) and Shishaldin (Alaska) [START_REF] Rasmussen | When does eruption run-up begin? Multidisciplinary insight from the 1999 eruption of Shishaldin volcano[END_REF]Singer et al., 2016), La Soufrière (Guadeloupe ; Lesser Antilles subduction zone) [START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF], Kelud (Indonesia; Sunda Arc subduction) (Utami et al., 2021) and Llaima (Chile) (Ruth et al., 2018) (Figure VII.1). Compared to the subduction rates, the Lesser Antilles is slower compared to other subduction zones studied here: New Zealand, Kamchatka and Sunda arc ones (~2 cm/yr [START_REF] Walker | Tarawera 1886, New Zealand -A basaltic plinian fissure eruption[END_REF] compared to ~4-5 cm/yr [START_REF] Wallace | Balancing the plate motion budget in the South Island, New Zealand using GPS, geological and seismological data[END_REF] and ~8 cm/yr (Gorbatov et al., 1997;[START_REF] Konstantinovskaia | Arc-continent collision and subduction reversal in the Cenozoic evolution of the Northwest Pacific: an example from Kamchatka (NE Russia)[END_REF]Minster and Jordan, 1978), respectively) or ~7 cm/yr for Indonesia, Southern Andes or Alaska [START_REF] Carver | Paleoseismicity and Neotectonics of the Aleutian Subduction Zone-An Overview[END_REF][START_REF] Göllner | Kinematic partitioning in the Southern Andes (39° S-46° S) inferred from lineament analysis and reassessment of exhumation rates[END_REF][START_REF] Tregoning | First geodetic measurement of convergence across the Java Trench[END_REF]. The difference in the subduction rates does not seem to have an effect on the timescales, as New Zealand volcanoes, the Lesser Antilles ones or volcanoes from the Andes both show short and long timescales (Figure VII.1). 43 % of the rest of the eruptions compiled are represented by evolved magmas (mostly arc volcanoes), with timescales ranging from 1 year to 1200 years (median of the datasets) (Figure VII.1). Dominica timescales plot in the part of the more evolved systems, dacitic to rhyolitic or phonolitic, on the order of several years to several decades (from 2 years for PPR3 to 34 years for Grand Fond (medians)) [START_REF] Ostorero | Time -window into the transcrustal plumbing system dynamics of Dominica (Lesser Antilles)[END_REF][START_REF] Solaro | A System Dynamics Approach to Understanding the deep Magma Plumbing System Beneath Dominica (Lesser Antilles)[END_REF] Systems with evolved magma compositions observed in caldera forming systems involve extremely high magma volumes (600 km 3 >10 km 3 or 40-60 km 3 of magma, respectively) but they generally have long timescales, of several decades (Bishop Tuff and Long Valley caldera (USA), Yellowstone or the 22-ka and 1600 BC (Minoan) caldera-forming eruptions of Santorini (Greece) [START_REF] Brugman | Common assumptions and methods yield overestimated diffusive timescales, as exemplified in a Yellowstone post-caldera lava[END_REF]Chamberlain et al., 2014;Druitt et al., 2012;Fabbro et al., 2018;Flaherty et al., 2018;Till et al., Opx from the Plinian phase of Bezymianny 1956 give more or less the same timescales as La Soufrière 5680 and 1657 (Guadeloupe), where a short repose time suggests a magma intrusion in the magmatic system, which is hotter than the mush (Metcalfe et al., 2021) (Figure VII.1). Opx from the blast are close to Ruapehu 1991 or 1971, where there were magma perturbations in the mush, 3 to 5 months before the eruptions [START_REF] Kilgour | Timescales of magmatic processes at Ruapehu volcano from diffusion chronometry and their comparison to monitoring data[END_REF]. Opx from the andesites of Kizimen are close from Ruapehu 1969 eruption and Bezymianny September 2012 eruption, where a magmatic injection occurred prior eruption, while the dacites give the same timescales as Bardarbunga, with pre-eruptive mixing and mush disaggregation was evidenced or from La Soufrière 1010 or Bezymianny 2008 (magmatic injection prior eruption) (Davydova et al., 2018;[START_REF] Kilgour | Timescales of magmatic processes at Ruapehu volcano from diffusion chronometry and their comparison to monitoring data[END_REF][START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF][START_REF] Mutch | DFENS: Diffusion Chronometry Using Finite Elements and Nested Sampling[END_REF] (Figure VII.1). Dominica timescales are close from evolved systems where magma recharge occurred prior to the eruptions (Merapi 2010 or Long Valley; Irazú) (Chamberlain et al., 2014;Costa et al., 2013;Ruprecht and Plank, 2013) or magma mixing processes (Santorini) (Fabbro et al., 2018) Grand Fond eruption gives the longer timescales, close to what was inferred for the Young Toba Tuff eruption or for Santorini 1600 BC, with recharge events that occurred a century before the eruption and mixing prior to the eruption (Druitt et al., 2012;Matthews et al., 2012) (Figure VII.1). Kizimen and Bezymianny timescales and magma dynamics then seem to be comparable to La Soufrière (Guadeloupe) or Ruapehu volcanoes, even if these timescales are calculated for small Plinian eruptions or phreatomagmatic eruptions for Ruapehu [START_REF] Kilgour | Timescales of magmatic processes at Ruapehu volcano from diffusion chronometry and their comparison to monitoring data[END_REF]. Thanks to Figure VII.1, we can see that pre-eruptive timescales of only 10 eruptions have been looked at by combining several crystal chronometers (including our studies on Kizimen and Bezymianny) (12% of the eruptions): Laacher See 12.9 ka [START_REF] Humphreys | Rapid pre-eruptive mush reorganisation and atmospheric volatile emissions from the 12.9 ka Laacher See eruption, determined using apatite[END_REF]Rout and Wörner, 2018;[START_REF] Sundermeyer | [END_REF], Long Valley (Chamberlain et al., 2014), Irazù [START_REF] Boyce | Apatite as a monitor of late-stage magmatic processes at Volcán Irazú, Costa Rica[END_REF][START_REF] Oeser | Combined Fe-Mg chemical and isotopic zoning in olivine constraining magma mixing-to-eruption timescales for the continental arc volcano Irazú (Costa Rica) and Cr diffusion in olivine[END_REF]Ruprecht and Plank, 2013), Santorini (Druitt et al., 2012;Fabbro et al., 2018;Flaherty et al., 2018), Nevado Del Toluca (Weber et al., 2019) and Yellowstone 262 ka [START_REF] Brugman | Common assumptions and methods yield overestimated diffusive timescales, as exemplified in a Yellowstone post-caldera lava[END_REF]Till et al., 2015). The timescales given by the different minerals are consistent, even if plag generally give longer timescales than the other crystals and apatite and mgt give shorter ones (Figure VII.1). This highlights the importance of confronting several minerals and their timescales.

To conclude, evolved systems give longer timescales than less evolved ones. Our studies will add some information on the relationships between the timescales recorded by several types of crystals and their combination to make complete models of the pre-eruptive magma dynamics. VII.1.2.2. Monitoring data 32 eruptions over the 86 eruptions (37 %) in Figure VII.1 compare petrological timescales with monitoring data, which shows the importance of monitoring data to try to forecast future unrest behaviors. The relationships between these two types of information were rather consistent, as shown in Kizimen 2010 eruption (Albert et al., 2019(Albert et al., , 2016;;Hartley et al., 2016;[START_REF] Kahl | Compositionally zoned crystals and real-time degassing data reveal changes in magma transfer dynamics during the 2006 summit eruptive episodes of Mt. Etna[END_REF]Kahl et al., , 2011;;Li et al., 2020;[START_REF] Magee | The Lead-up to Mount Etna's Most Destructive Historic Eruption[END_REF][START_REF] Marti | Correlation of Magma Evolution and Geophysical Monitoring during the 2011-2012 El Hierro (Canary Islands) Submarine Eruption[END_REF][START_REF] Moussallam | Fast ascent rate during the 2017-2018 Plinian eruption of Ambae (Aoba) volcano: a petrological investigation[END_REF]Pankhurst et al., 2018;Rae et al., 2016;[START_REF] Rasmussen | When does eruption run-up begin? Multidisciplinary insight from the 1999 eruption of Shishaldin volcano[END_REF]Ruprecht and Plank, 2013;Ruth et al., 2018;Sundermeyer et al., 2020a), apart from some eruptions where only syn-eruptive monitoring data were compared or where seismic signals were not of high quality or altered by a seismic activity on the specific volcanoes (Costa et al., 2013;Davydova et al., 2018;[START_REF] Kilgour | Timescales of magmatic processes at Ruapehu volcano from diffusion chronometry and their comparison to monitoring data[END_REF]Saunders et al., 2012;Utami et al., 2021). However, as the probability of an eruption of a magnitude 7 or greater in this century is 1 in 6 [START_REF] Cassidy | Huge volcanic eruptions: time to prepare[END_REF][START_REF] Lin | Magnitude, frequency and climate forcing of global volcanism during the last glacial period as seen in Greenland and Antarctic ice cores (60-9 ka)[END_REF], monitoring efforts have to be made regarding the observations of active volcanoes. Only approximately 27% of the eruptions since 1950 have been monitored with at least one instrument (seismometer), and remote sensing monitoring can be useful for volcanoes where ground-based monitoring is not possible [START_REF] Cassidy | Huge volcanic eruptions: time to prepare[END_REF]. However, current satellites lack the necessary resolution in time and space to observe volcanoes in real time, indeed, after the Tongan eruption in 2022, changes at the volcano were only seen by satellite radar images 12 hours after the eruption [START_REF] Cassidy | Huge volcanic eruptions: time to prepare[END_REF]. VII.1.2.3. Experimental studies Chaotic magma mixing experiments of a synthetic basalt and a natural basaltic sample, as well as two natural samples (basaltic andesite and dacite) were carried out to constrain both the timescales of the mixing event that led to the formation of the andesite in the reservoir and the mingling event prior to the 2010 Kizimen eruption (Chapters VI.3-4). As crystallization controls the chemical and textural evolution of magmas at pre-eruptive conditions, they are important processes to constrain during mixing, as well as the kinetics of mixing. Our studies allowed to investigate these two processes together, whereas they were generally investigated separately [START_REF] De Campos | Enhancement of magma mixing efficiency by chaotic dynamics: an experimental study[END_REF], 2008, 2004;[START_REF] Erdmann | Textures of Peritectic Crystals as Guides to Reactive Minerals in Magmatic Systems: New Insights from Melting Experiments[END_REF][START_REF] Hammer | An experimental study of the kinetics of decompression-induced crystallization in silicic melt[END_REF][START_REF] Laumonier | Experimental simulation of magma mixing at high pressure[END_REF][START_REF] Mollard | Decompression-induced Crystallization in Hydrated Silica-rich Melts: Empirical Models of Experimental Plagioclase Nucleation and Growth Kinetics[END_REF]Morgavi et al., 2013bMorgavi et al., , 2013a;;Perugini et al., 2015aPerugini et al., , 2010;;Perugini and Poli, 2012;[START_REF] Shea | Kinetics of cooling-and decompression-induced crystallization in hydrous mafic-intermediate magmas[END_REF].

Residence times of the microlites were rather short in the natural banded andesites, ranging from minutes to hours. Compared to other magma mixing experimental studies and geochemical numerical modelling, the magma mingling timescales are in the same order of magnitude, of the orders of hours (with recent simulations showing that complete hybridization of magmas for the Campi Flegrei magmatic systems) (Montagna et al., 2015;[START_REF] Morgavi | Decrypting Magma Mixing in Igneous Systems[END_REF]Morgavi et al., , 2017;;Perugini et al., 2010). The timescales found here are also close from the mixing-to-eruption timescale in the order of a few minutes for Campi Flegrei volcanic system, seen from magma mixing experiments and statistical analysis (Perugini et al., 2015a). These timescales are far shorter than the timescales inferred with chronometry studies, seen in the section before VII. 1 showing that these magma mixing/mingling phenomena can participate in the triggering of eruptions on a very little time window (Figure VII.1).

VII.2. Reflections on the method strategies

Reflecting on the method of timescales modelling, several limitations are linked to these estimations, such as the fact that the crystals were not always suitable for modelling (possibility of being affected by growth processes) or uncertainties associated to these timescales due to the temperature estimates and the choice of the diffusion coefficient (uncertainties associated with their determinations). A strategy to decipher the timescales in crystals, for example in opx could consist of continuing to mount a lot of crystals and always acquiring SEM images of the zoned crystals prior to the EPMA analysis to check their greyscale content. This evaluation of the greyscale data could ensure that the En content difference of the zoned opx is sufficient to be modelled or that their profiles are typical of diffusion (by using the ImageJ software). Then, for the opx model of the timescales, it could be of interest to always test several temperatures to grasp the importance of a good estimation of this parameter, as well as the effect of the oxygen fugacity, as it has been done in Chapter V. I wanted to specifically tackle these aspects, as throughout my thesis, these comments on the variations of these parameters and the meaning of the timescales were evoked.

To avoid the bias of the AUTODIFF spreadsheet model (Allan et al., 2013;Couperthwaite et al., 2020) which is due to the fitting of the profiles by eye, it would be useful to use the code developed in [START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF] which uses a misfit function to fit the timescales and statistical elements to compare the timescales. Modifications of this code could be done to integrate an automatic fitting between EPMA data with greyscale data of zoned profiles, as the current code only uses the greyscale data, which represents a bias. The implementation of the diffusion coefficient of Ganguly and Tazzoli (1994) could also be done in the code, as this coefficient is the only one determined in opx with a similar En content as the opx of our datasets.

Concerning the diffusion coefficients, as seen in the Chapters IV-VI.1, Ganguly and Tazzoli (1994) did not constrain the dependence on the oxygen fugacity fO2 but hypothesized its dependence by analogy with olivine. We are aware that not considering the fO2, as it was just hypothesized by Ganguly and Tazzoli (1994) can lead to an overestimation of the timescales. We decided not to use another diffusion coefficient expression, as the interdiffusion coefficient determined by experimental annealing of diffusion couples and given by Dohmen et al. (2016) in Chapters IV-VI.1, were not adapted to the En content of the opx in this thesis (established for high En contents). Using this coefficient could lead to erroneous timescales due to compositional effects, even if it took the fO2 into account. A new experimental study using the En content of the volcanic systems studied here and oxygen fugacity ranges would be necessary to obtain a new diffusion coefficient that would be appropriate.

Finally, to ensure having a complete model of the pre-eruptive dynamics, the storage conditions of the magmas such as temperature, pressure (depth) or oxygen fugacity should also be well-constrained. Geothermometry investigations based on Fe-Ti oxides and mineral-liquid thermobarometry should be done, as well as measures of major elements and volatile content in melt inclusions and of the residual glasses near crystals which should be added in the strategy, if no literature data on the same samples is available. For Kizimen and Bezymianny, despite a high number of opx and other crystals mounted (4301 opx, plag, mgt and amph for Kizimen and 3827 for the blast and Plinian phase of Bezymianny), the melt inclusions identified in the crystals were rare and when present, were generally too small (less than 10 microns) so EPMA or SIMS analyses were not always possible. For Kizimen and the blast and Plinian phase of Bezymianny, only Fe-Ti oxide geothermometry investigation were then carried out. The magma at the origin of the 1956 Plinian phase of Bezymianny climactic activity have storage conditions that are better constrained (melt inclusions from d'Augustin (2021)) so these ones were used for comparison, as well as Auer et al. (2018) conditions for Kizimen 2010 eruption.

Concerning the experimental studies, a new experimental strategy could consist in the measure of the oxygen fugacity during the experiment to know more precisely the liquidus temperature of the crystals thanks to MELTS simulations. Despite the simplifications carried out in the experiments, such as the absence of volatiles in the melts (as powders were melted under ambient pressure), the rheology of the melts involved were not the same, as well as the compositions of the endmembers, overall, a comparison with the natural system of Kizimen can be carried out.

Finally, one of the aims of my thesis was to make a step towards an early-warning system based on the timescales and pre-eruptive dynamics deciphered. This may have been achieved with Kizimen 2010 eruption, as the correlation between petrological timescales and geophysical monitoring signals is shown. A systematic pattern in the pre-eruptive timescales was identified for 5 past Plinian-type explosive eruptions in the last 5000 years of Montagne Pelée (Martinique), showing that it is possible to have this systematic pattern for unmonitored eruptions in a storage system that is stable (Boudon et al., 2018). Indeed, it was shown that the time between changes in magma storage conditions and eruptions mainly occurred in the last 2 years (Boudon et al., 2018). To be able to identify this kind of systematic pattern in the timescales through several eruptions for a specific volcano, we would need to study more eruptions of Kizimen. Bezymianny climactic phase in 1956 study was not focused on the correlation with monitoring data but on determining timescales after 1000 years of repose time and to constrain the pre-eruptive dynamics of this specific eruption. As this eruption was unique, it would be difficult to carry out the same study on the following eruptions to identify a systematic pattern in the timescales as they have different eruptive dynamics (open conduit) and involve more mafic magmas. The study of Davydova et al. (2018Davydova et al. ( , 2021) ) represents a step forward for early-warning systems, as several eruptions (from 2006-2012 and the 2017 eruption of Bezymianny) were studied and showed that the magmatic reservoirs were fed by basic magmas and that magma mixing occurred prior to these eruptions.

Several steps and evolution are needed to obtain an early warning system, as it is more than a mechanism to issue warnings of impending hazard events (United Nations, 2015). To become a system, it has to be constituted by four integrated sub-systems: a warning sub-system where hazards are monitored and where forecasts and warnings issued; a risk information sub-system in which risk scenarios are generated for areas and population likely to be affected; a preparedness sub-system which indicates the actions that should be taken to reduce loss and damage and finally, by a communication sub-system that allows a communication of information on hazard events, risks and appropriate preparedness strategies to those at risk (United Nations, 2015). The importance of early warning systems and communications between volcanologists, public officials and the public was particularly evidenced following the eruption of La Soufrière in St Vincent in 2021, when volcano-tectonic earthquakes increased and evacuation of some areas began, when the eruptive behavior shifted from extrusive to explosive in April 2021 [START_REF] Graham | Managing stakeholder relationships for improved situation awareness during volcanic emergencies: An Eastern Caribbean case study[END_REF].

VII.3. Perspectives

In this thesis, we expanded the timescales dataset for seven eruptions of three different arc volcanoes in the world and improved our understanding of pre-eruptive magma dynamics before high VEI eruptions in the past. We also added some melt inclusions and residual glasses data for these volcanoes, that can improve the geochemical classifications of these eruptions.

Open questions are still unanswered, about the organization of the magma reservoirs throughout the mush, needing interdisciplinary studies to have a clearer picture; about the reactivation of these systems and what determines their evolution towards eruptible states, with the triggering event and ascent timescales or about the pre-eruptive dynamics and the diffusive timescales recorded by different crystals.

One perspective of this work would be to perform diffusive experiments to determine a new interdiffusion coefficient for opx in the compositional range from En40 to En80 (for example), by constraining the fO2 dependence. These experiments would allow us and several authors to reevaluate the timescales modelled in opx using Dohmen et al. (2016), if their opx compositional ranges were not the same, or the ones using Ganguly and Tazzoli (1994) without the fO2 (even if the oxygen dependence has been shown to be weak) (Allan et al., 2017;Dohmen et al., 2016).

Diffusion in other minerals of the same eruptions could also be systematically combined. Modelling diffusion timescales on zoned plagioclase or in magnetites, as well as in clinopyroxenes or apatites, when present, and interpreting them together could be compared, as diffusion chronometry conducted with multiple elements in multiple minerals produces accurate temporal results [START_REF] Brugman | Common assumptions and methods yield overestimated diffusive timescales, as exemplified in a Yellowstone post-caldera lava[END_REF]. Other softwares such as NIDIS for clinopyroxenes (Petrone et al., 2018(Petrone et al., , 2016;;[START_REF] Rout | Long-Term Temperature Cycling in a Shallow Magma Reservoir: Insights from Sanidine Megacrysts at Taápaca Volcano[END_REF][START_REF] Rout | Constraints on non-isothermal diffusion modeling: An experimental analysis and error assessment using halogen diffusion in melts[END_REF] or DIPRA software for olivines [START_REF] Girona | DIPRA: A user-friendly program to model multi-element diffusion in olivine with applications to timescales of magmatic processes[END_REF]Lynn et al., 2017;[START_REF] Moussallam | Fast ascent rate during the 2017-2018 Plinian eruption of Ambae (Aoba) volcano: a petrological investigation[END_REF]Sundermeyer et al., 2020a) could be used. The new software Diffuser [START_REF] Wu | Diffuser: A user-friendly program for diffusion chronometry with robust uncertainty estimation[END_REF] could provide the modelling of the diffusion of different elements in orthopyroxenes, such as rare earth elements [START_REF] Cherniak | Rare earth element diffusion in natural enstatite[END_REF][START_REF] Rutherford | Neodymium diffusion in orthopyroxene: Experimental studies and applications to geological and planetary problems[END_REF] or Ti, widely used to study magma differentiation processes and quantify cooling rate [START_REF] Cherniak | Ti diffusion in natural pyroxene[END_REF] or fast diffusion with H (Demers- Roberge et al., 2021;[START_REF] Stalder | Hydrogen self-diffusion in pure and Cr-doped enstatite[END_REF]. Isotopic zoning analysis of the crystals could also be conducted, as recent studies in olivines showed that by combining elemental and isotopic zoning studies in individual minerals, it allowed to refine the timing of magmatic eruptions and constrain the origin of the zoning patterns [START_REF] Oeser | Combined Fe-Mg chemical and isotopic zoning in olivine constraining magma mixing-to-eruption timescales for the continental arc volcano Irazú (Costa Rica) and Cr diffusion in olivine[END_REF][START_REF] Oeser | Processes and time scales of magmatic evolution as revealed by Fe-Mg chemical and isotopic zoning in natural olivines[END_REF].

In order to have more insights in the pre-eruptive dynamics of Bezymianny volcano, study of recent eruptions after the laterally directed blast could help understanding the magma dynamics. A study on a lava dome fragment of Bezymianny was begun (Bezy 15) and we observed reverse zoned opx, which are archives of a heating process prior eruption. A more complete study could be continued, as the crystals mounts of opx, plag, amph and mgt are already done. A study of March 2019 clasts could be performed to compare the pre-eruptive timescales with ones found by Davydova (2018[START_REF] Davydova | Petrological evidence of rapid evolution of the magma plumbing system of Bezymianny volcano in Kamchatka before the December 20th, 2017 eruption[END_REF], to model the timescales prior to recent eruptions of Bezymianny and compare them with the other pre-eruptive magma dynamics of the eruptions that followed the one in 1956.

Other volcanoes from Kamchatka or in the Lesser Antilles could be studied in the future (even if recent and on-going diffusion studies focus on La Soufrière (Guadeloupe) and Montagne Pelée (Martinique) (Boudon et al., 2018;[START_REF] Metcalfe | Insights into the Origin and Dynamics of Past Eruptions of La Soufrière de Guadeloupe from Melt Inclusions: Implications for Future Unrest and Environmental Impact[END_REF][START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF]), but it would be interesting to look at crystals from St. Vincent 2021 eruption, for example or Shiveluch volcano. It would of interest to carry out more petrological monitoring studies and compare the timescales found in evolved systems with a basaltic volcano of Kamchatka or other with subduction zones volcanoes, in Indonesia, South America or Japan where they are a threat to communities… It would also be interesting to look at effusive eruptions, in other volcanic contexts, such as Mayotte volcano and other volcanic features [START_REF] Berthod | The 2018-ongoing Mayotte submarine eruption: Magma migration imaged by petrological monitoring[END_REF][START_REF] Feuillet | Birth of a large volcanic edifice offshore Mayotte via lithosphere-scale dyke intrusion[END_REF] to delve into the formation of volcanoes and their plumbing system architectures.

Another perspective could consist in focusing on more shallower magma dynamics and ascent phenomenon, that will greatly influence the eruptive style and the volcanic hazards. As my PhD mainly focused on pre-eruptive processes in the magmatic reservoir, it would be of interest to study these preeruptive processes during ascent in the magma conduit to understand better the eruptive behavior and phases. A better understanding of eruptive styles in evolved systems would be allowed by interdisciplinary efforts combining petrology, geochemistry and monitoring of volcanoes to investigate the factors influencing eruption dynamics. By studying the matrix textures and microlites, it will allow to refine the composition of the original melt from which the microlites grew, to understand the shallow storage regions and their depths. Decompression experiments could also permit to constrain ascent rates. This would permit to focus more on the timescales of magma ascent and their dynamics, degassing processes, through textural studies of the experimental samples (Arzilli et al., 2019a;[START_REF] Cashman | Groundmass crystallization of Mount St. Helens dacite, 1980?1986: a tool for interpreting shallow magmatic processes[END_REF][START_REF] Martel | Decompression experiments as an insight into ascent rates of silicic magmas[END_REF]. Finally, constraining the factors that control the eruptive style, which are the magma rheology (viscosity), then volatile content and crystals, as well as magma ascent rates can bring precious information to mitigate volcanic crisis.

Finally, studying the correlations between recent and well-monitored eruptions with monitoring signals was the first objective of the V-Care ANR project in which my PhD is integrated. To move from these diffusion studies to an early-warning clock, several eruptions from the same volcano have to be studied to see if different parts of the system were reactivated between the eruptions and if they record the same history. It is also important to work on the significance of the timescales obtained by diffusion in zoned crystals. This study of the presence or not of a systematic pattern of the pre-eruptive dynamics is really necessary to then apply this pattern of magmatic reactivation to other subduction zones volcanoes that did not produce historical magmatic eruptions. These eruptive scenarios have then to be considered in the sub-systems evoked in VII.2. Future modelling of the eruptive scenarios, considering the magma dynamics evidenced prior to these eruptions, have to be conducted, with different responses regarding the unrest signals that would potentially be detected. It has to be combined with the second objective of the V-Care ANR project, which is on the societal impacts and on the reactions of the populations for an early awareness of the eruptions. Integrating these different scenarios for an early-warning system could allow to better prepare the world for massive eruptions, as the probability of an eruption of magnitude of 7 or greater is of 1/6 for this century [START_REF] Cassidy | Huge volcanic eruptions: time to prepare[END_REF].

INTRODUCTION

How magma is stored at depth and how magma bodies assemble prior to eruption are two questions that are still widely debated. The architecture and associated dynamics of magma plumbing systems for silicic volcanoes have been in the last decade envisaged as transcrustal magma systems (Cashman et al., 2017)f o l l o w i n g the pioneering work of Bachmann and Bergantz (2003).

In the process of sequential accretion of sills by which large crustal magma reservoirs form (Annen, 2009;Bachmann and Huber, 2016;[START_REF] Blundy | Crustal magmatic systems from the perspective of heat transfer[END_REF], the resident magma can cool and partially crystallize in between successive magma pulses to become a crystal mush (Marsh, 2006). Crystallinity stays in the 50-60% range and the solidus is not reached quickly because of the release of latent heat during crystallization (Marsh, 1989). This type of highly-crystallized crustal reservoir raises questions regarding magma mobility and the timescale over which magma can be remobilized through the rejuvenation of the mush to become eruptible melt. Several mechanisms have been proposed to enable mush rejuvenation and possible melt extraction: selfmixing (Couch et al., 2001), gas sparging (Bachmann and Bergantz, 2003;Bachmann and Huber, 2016;Huber et al., 2010), or convective stirring (Huber et al., 2009). In any case, underplating mafic magmas bringing heat and/or volatiles are considered as major players in the reactivation of mush. Heat and volatile transfers change the rheological properties of the mush (buoyancy, crystallinity), favoring its mobility (formation of melt connections, volatile foams). The efficiency of the enthalpy (±mass) exchange with newly injected recharge magma of higher enthalpy necessary to unlock a crystal mush highly depends not only on the size but also on the composition and temperature of the intruding magma body. Estimated timescales suggest that the reactivation timescale is much longer than the following homogenization timescale (Huber et al., 2012).

Detailed petrological investigations of the compositions of crystals, their melt inclusions and the residual melt provide strong constraints on the dynamics of the magmatic plumbing system in terms of architecture and timescales. Although such studies are major undertakings, they are highly informative (Ruth et al., 2018;Albert et al., 2019). Using the method developed by Palm (2005), the pre-eruptive dynamics of recent eruptions of Mt. Etna has been explored using the sequential variations in olivine crystal compositions, represented schematically in a box model and called Crystal System Analysis (CSA) (Kahl et al., 2011;Kahl et al., 2015).

Coupling CSA with diffusion timescales can be used to show variable evolution of magmatic interactions prior to eruption and showed variable magmatic pathway systematics between eruptions. In recent decades, intracrystalline diffusion modeling has been increasingly used to study timescales of pre-eruptive processes, built on the pioneering works of Costa et al. (2003), Costa and Chakraborty (2004) and Chakraborty (2008). Because of the various diffusion ratesofelementsincrystals (Costa et al., 2020), this technique allows investigation of timescales spanning minutes [START_REF] Charlier | Lithium concentration gradients in feldspar and quartz record the final minutes of magma ascent in an explosive supereruption[END_REF] to thousands of years (Allan et al., 2017) and, being applicable to single crystals, it can yield timescales of specific processes. Several studies, covering various volcanic settings and involving different compositions of magmas and eruptive dynamics have already been conducted with different element/crystal pairs, including olivine (Costa and Chakraborty, 2004;Costa and Morgan, 2010;Kahl et al., 2011;Kahl et al., 2015), plagioclase [START_REF] Zellmer | Magma emplacement and remobilization timescales beneath Montserrat: insights from Sr and Ba zonation in plagioclase phenocrysts[END_REF][START_REF] Zellmer | Magma emplacement and remobilization timescales beneath Montserrat: insights from Sr and Ba zonation in plagioclase phenocrysts[END_REF]Ginibre et al., 2004;Ruprecht and Wörner, 2007;[START_REF] Saunders | The magmatic evolution of the Whakamaru supereruption, New Zealand, constrained by a microanalytical study of plagioclase and quartz[END_REF]Shcherbakov et al., 2011;Druitt et al., 2012), pyroxenes (Morgan et al., 2004;Allan et al., 2013;Barker et al., 2016), magnetite (Boudon et al., 2015)andquartz (Matthews et al., 2012;Chamberlain et al., 2014).

In the Lesser Antilles Arc, the volume of magma involved in most of the eruptions is relatively low along the volcanic arc (generally <0.5 km 3 ) as evidenced by many historical and prehistorical eruptions (Lindsay et al., 2005). Dominica is an exception: it is the sole island of the arc with several active volcanoes. Together with the neighboring islands of Guadeloupe and Martinique, it concentrates half of the active volcanoes of the arc within its central part, only accounting for 1/5 of its total length (Figure 1A). The pioneer works on the volcanic activity of Dominica recognized a voluminous silicic eruption, dated between 50 and 30 ka and called "the Roseau Tuff" [START_REF] Sigurdsson | Partly--welded pyroclast flow deposits in Dominica, Lesser Antilles[END_REF][START_REF] Carey | The Roseau ash: deep-sea tephra deposits from a major eruption on Dominica, Lesser Antilles arc[END_REF]. Approximatively 58 km 3 Dense Rock Equivalent (DRE) of magma has been first estimated for this eruption, making it as the most voluminous eruption of the arc. Though recent works now recognized several eruptions, in particular in the central part of Dominica (Lindsay et al., 2005;Smith et al., 2013;Boudon et al., 2017), with necessarily lower emitted volumes, these volumes still contrast with those recognized for the other active volcanoes. For instance, the ongoing eruption of Soufrière Hill's of Montserrat only emitted 1 km 3 of magma (Wadge et al., 2014) and the famous devastating eruption of Montagne Pelée (Martinique) in 1902-1905 was responsible of the emission of only 0.2-0.3 km 3 (Boudon, 1983). Dominica volcanic activity thus offers the unique opportunity to investigate the magma dynamics in the plumbing system prior to repeated large silicic eruptions, only recognized on this island.

Here, the aim is to identify and characterize the pre-eruptive disrupting events leading to two successive large silicic eruptions (∼3-5km 3 DRE; Boudon et al., 2017) in the central part of Dominica, Layou (∼51 ka) and Roseau (∼33 ka), and their associated timescales. For that purpose, we present an original study using orthopyroxenes that combine a variant of the Crystal System Analysis method and intracrystalline diffusion modeling. More precisely, we performed a detailed petrological investigation of pumices ejected during the first eruptive Plinian phase of these two major eruptions. We focus our analysis on orthopyroxenes, which crystallize over a wide range of temperature, pressure and oxygen fugacity. Thus, orthopyroxene crystals may be good tracers of evolving crustal storage conditions. This work builds upon a revised chronostratigraphy of the eruptions (Boudon et al., 2017), experimental investigations of the pre-eruptive magma storage conditions (Solaro et al., 2019), and a detailed study of major, trace, and volatile element compositions of melt inclusions (Balcone-Boissard et al., 2018). The results are discussed for their implications in terms of the more general framework of the magma plumbing system architecture and dynamics beneath Dominica. Furthermore, we assess the rates of silicic magma rejuvenation prior to voluminous explosive volcanic eruptions.

GEOLOGICAL CONTEXT AND PREVIOUS WORK Geological Context

The Lesser Antilles Arc results from the westward subduction of the North-South American plates beneath the Caribbean plate, at an average rate of 2 cm/yr (Wadge, 1984). The volcanic arc is 850 km long, and extends from the island of Saba in the north to Grenada in the south, with an eastward convexity. Dominica is located in the central part of the arc (Figure 1).

North of Dominica, the arc splits into two branches: the Outer or Older arc (Eocene to mid-Oligocene) to the north-east and an Inner or Recent arc (Miocene to Recent) trending north-west. This jump in the northern sector has been ascribed to the ancient subduction of an aseismic ridge in the center of the arc causing a rupture in the subducting plate and a displacement of volcanic activity to the west in the early-Miocene (∼20 Ma; Bouysse and Westercamp, 1990). South of Dominica, the two branches merge, creating a single volcanic arc active since the Miocene. The active parts of the arc (Saba to Dominica in the North, and Dominica to Grenada in the South) have twelve, currently-active volcanic centers, the best-known being Soufrière (Guadeloupe), Montagne Pelée (Martinique), Soufrière Hills (Montserrat) and Soufrière (St.-Vincent). Conversely to the other islands, Dominica is characterized by four, potentially-active volcanic centers, from North to South: Morne aux Diables, Morne Diablotins, Morne Trois Pitons-Micotrin and Morne Plat Pays Volcanic Complex (Lindsay et al., 2005; Figure 1B).

Five ignimbritic eruptions are currently recognized on Dominica: Grand Bay, Roseau, Grand Fond, Layou and Grande Savane ignimbrites (Lindsay et al., 2005;Smith et al., 2013;[START_REF] Howe | Reevaluation of the Roseau Tuff eruptive sequence and other ignimbrites in Dominica, Lesser Antilles[END_REF]Boudon et al., 2017; Figure 1B). These deposits have been recently re-investigated by Boudon et al. (2017) who derive new 14 C ages for the Layou, Roseau and Grand Fond ignimbrites respectively at ∼51, ∼33 and ∼24 ka; each eruption involved magma volumes of the order of 3-5km 3 DRE. On the basis of geological arguments and different geochemical signatures observed in Roseau magmas compared to Layou ones, the Roseau ignimbrite originates from Morne Trois Pitons-Micotrin. The Layou ignimbrite likely originated from the Morne Diablotins eruptive center. The deposit of each eruption is subdivided into a basal Plinian fallout (<10% of total erupted volumes) (Figures 1C,D), overlain by thick (up to 100 m) pumiceous pyroclastic density current deposits that vary in grain size, welding and thickness distribution (Figures 1E,F).

Previous Works: Petrology of Pumice Clasts and Pre-eruptive Magma Storage Conditions Mineralogy Whole rock, melt inclusions, and residual glass compositions (major and trace elements) have been studied together with the crystal contents of the natural samples of the Roseau and Layou eruptions (Boudon et al., 2017;Balcone-Boissard et al., 2018). Magma of the Roseau eruption is intermediate between andesite and dacite (62-63 wt% SiO 2 , ∼4.5-5.0 wt% Na 2 O+K 2 O) whereas magma from the Layou eruption is dacitic (>64 wt% SiO 2 , ∼5wt% Na 2 O+K 2 O). For both eruption deposits, the residual glass is rhyolitic (76.5-78.0 wt% SiO 2 ,5 . 5 -7.0 wt% Na 2 O+K 2 O). The samples are porphyritic with ∼30 ± 5 vol% phenocrysts (Boudon et al., 2017) consisting of plagioclase, orthopyroxene, amphibole, clinopyroxene, Ti-magnetite and ilmenite embedded in a microlite-free residual glass (Solaro et al., 2019). Modal composition of the magma of each eruption shows that plagioclase is the most abundant phenocryst (70%), followed by orthopyroxene (15-17%), amphibole (up to 17% for Layou, though <1% for Roseau) and, in lesser proportions, clinopyroxene and Fe-Ti oxides. Rare quartz phenocrysts are also recognized in the Layou deposit. Plagioclase crystals show oscillatory zoning patterns and resorption textures, with anorthite content (An, expressed in mol %) spanning a broad range between An 42 to An 90 (Solaro et al., 2019), with some high An spikes. Amphiboles in Layou are ferrimagnesio-hornblende with Mg# [Mg/(Mg + Fetot)] ranging from 0.52 to 0.60. Clinopyroxenes are augites with Mg# ranging from 0.48 to 0.71. Fe-Ti oxides are titanomagnetites (Mag 71-75 ;M a g magnetite mol%) and ilmenites (Ilm 87 ;Ilm ilmenite, mol%). The orthopyroxene compositions, the focus of this study, are presented hereafter.

Melt Inclusions

Only glassy Melt Inclusions (MIs) with no post-entrapment modifications have been investigated, as detailed in Balcone-Boissard et al. (2018). MIs are ubiquitous in orthopyroxenes (sizes up to ∼50 µm). In plagioclase, scarce MIs are smaller than 40 µm. The MIs are all rhyolitic regardless of their host crystal, with some variations in Na 2 O+K 2 O and SiO 2 contents. MIs are among the most water-rich yet recorded (up to 8 wt% H 2 O). Carbon dioxide contents are generally low (<650 ppm), although MIs trapped at high pressures contain higher CO 2 contents (up to 3,500-4,000 ppm). Some low-pressure (<100 MPa) inclusions have elevated CO 2 contents (up to 1,100-1,150 ppm), suggestive of fluxing of shallow magmas with CO 2 -rich fluids. CO 2 -trace element systematics indicate that melts were volatilesaturated at the time of entrapment (Balcone-Boissard et al., 2018).

Magma Storage Conditions

Phase equilibrium experiments starting from glasses of Roseau and Layou bulk compositions have been performed in an internally-heated pressure vessel, spanning a temperature range of ∼800-900 °C, a pressure range of 200-400 MPa (under water saturation and under-saturation), and fixed oxygen fugacity of ≈ NNO +0.6 (NNO Nickel-Nickel oxide buffer) (Solaro et al., 2019). By experimentally reproducing the natural melt and phenocryst rim compositions, phenocryst assemblages, phase proportions, and water contents in MIs, the magma storage conditions have been determined to be close to a pressure of ∼400 MPa, a temperature of 850 ± 25 °C with 7-8 wt% H 2 O dissolved in melt. These parameters will be used for the diffusion modeling below.

MATERIAL AND METHODS

Sample Types and Preparation

Studied pumices and crystals come from the field campaigns performed previously for chronostratigraphic, melt inclusions and experimental petrological studies (Boudon et al., 2017;Balcone-Boissard et al., 2018;Solaro et al., 2019). The samples consist of white, homogeneous pumice clasts collected from the basal Plinian fallout deposits of Layou and Roseau (Figures 1C,D compositional boundary conditions are the same, the difference between a profile at two different times is merely a stretch factor along the mean diffusion direction. This also holds true in the case of composition-dependent diffusion, where the curve shape is determined by the degree of compositional contrast, but the curve width is dependent on time (Couperthwaite et al., 2020). By using a library of profiles calculated for particular values of compositional contrast, the appropriate curve shape to apply to a natural crystal can be selected from the library based upon compositional data, and then scaled in width to solve for time, leading to very rapid processing. The scaling factor needed is a simple expression of crystal composition, the diffusivity used for the library profiles, the diffusivity in the crystal and the ratio of diffusion timescales in the library and the timescale of the crystal. As the scaling factor can be determined from data, the timescale for the crystal can be simply solved as the only unknown. The internal consistency of timescales obtained by this modeling method is tested by modeling several profiles on the same core-rim boundary of the same crystal and verifying for the acquisition of comparable timescales.

As we undertake modeling without correction for oxygen fugacity, the largest source of uncertainty in the diffusion timescales retrieved by this method is derived from the temperature. Varying the input values of temperature by ±25 °C in the Fe-Mg interdiffusion modeling for the same boundary implies change up to one order of magnitude in the calculated timescales (Allan et al., 2013). Here a Monte Carlo simulation was used to estimate the integrated time uncertainty from the different sources of uncertainty: final uncertainties on our reported timescales arise from uncertainties in diffusivity (D) calculations, which derive in turn from 1) uncertainties in measurements of diffusion coefficients (D 0 ,E a ); 2) uncertainty in temperature (the Arrhenius relation gives an inverse exponential dependence on diffusion); 3) considerations about the resolution of greyscale values of the BSE images used for the profile calibration (pixel size, 1σ uncertainty on pixel size, 1σ uncertainty on greyscale values defining a plateau). The resolution of BSE images has usually a small effect on final uncertainties (∼1%; Costa and Morgan, 2010). Uncertainties on temperature, the pre-exponential factor (D 0 ) and the activation energy (Ea) all have a logarithmic effect on final calculated timescales such that the final uncertainties are not Gaussian in time but asymmetric, with larger error bar on the long-time side and smaller error bar on the short-time side, with uncertainties approaching Gaussian behavior in log-time.

Diffusion Parameter

Here we have used the diffusion coefficient resulting from the parameterization of Ganguly and Tazzoli (1994), without correction for oxygen fugacity.

LogD Fe -Mg c-b -9.54 + 2.6 X Fe -12530

T ( 1
)
where D is the Fe-Mg interdiffusivity, D 0 -9.54 (m 2 s -1 ) is the pre-exponential factor, X Fe is the molar fraction of the ferrosilite component and T is the temperature in Kelvin. The Ganguly and Tazzoli (1994) equation is formulated for T 500-800 °C, oxygen fugacity of IW to IW + 0.8 (IW Iron-Wüstite buffer) and X Fe 0.10-0.50. The question of considering or not the oxygen fugacity is subject to controversy (Supplementary Table S1). In their paper, Ganguly and Tazzoli (1994) conjectured that the oxygen fugacity dependence could follow the same form as that for olivine, with an exponent of 1/6, assuming the same mechanism of vacancy generation through iron oxidation, but had no further data upon which to constrain that hypothesis. This parameterization appears to be consistent with the experimental results of Dohmen et al. (2016), who argue for a minor dependence of the diffusion coefficient on oxygen fugacity (n 1/20 instead of 1/6), although that study examined pyroxene very close to enstatite, and which is not a good match for these samples. A further study of Klugel (2001) in iron rich orthopyroxenes (in the temperature range 1,050-1,200 °Ca n d at QFM [quartz-fayalite-magnetite] oxygen buffer) reported a diffusivity of D Fe-Mg 3 × 10 -19 m 2 s -1 at T 1,130 °C. If extrapolated to higher temperatures, the results of Ganguly and Tazzoli (1994) are in good agreement with those of Klugel (2001).

In summary, all modeled profiles and timescales have been here calculated for a fixed temperature of 850 °C( ±25 °C) based upon the experimental work of Solaro et al. (2019) on these magma compositions. This yields diffusivities of the order of D ∼10 -20 m 2 s -1 (Ganguly and Tazzoli, 1994), without dependence of the diffusivity on oxygen fugacity. The initial condition is considered to be a step-wise profile (i.e., instantaneous growth at a fixed composition followed by diffusional relaxation of the interface), which yields maximum timescales if all broadening occurs via diffusion.

Compositional profiles for these eruptions have been interpreted as pure diffusion profiles with no or limited recorded growth, i.e., considering profiles as due purely to diffusion. This constrains maximum timescales, because if the profiles have a growth component, the required timescale is shortened. The fitted profiles have features typical of diffusive character-smooth compositional variation of sigmoid shape (Supplementary Figure 2). Oscillations can be seen toward the rim but are minor in amplitude (±0.2 mol% or less) and could plausibly be attributed to noise. Typically, growthdominant profiles would display linear trends or curves on graphs of composition vs. distance, which would not be congruent nor consistent with a diffusion sigmoid calculated for the given initial compositional contrast.

Crystal System Analysis Approach

The CSA technique has been successfully applied to the study of Etna, in order to define the main active magmatic environments (ME) prior eruption from crystal compositional zoning, their connections (by heating, decompression, mixing, mingling, crystal settling, etc) and the pre-eruptive processes occurring at depth (Kahl et al., 2011;Kahl et al., 2015). Following this definition, a ME is not necessarily a single batch of magma, but may be different part of the same reservoir at depth. In the case of Etna, a variety of ME were described based on olivine-melt equilibria and their occurrence with time. We applied this method to orthopyroxenes (opx) that show variations in major-element composition (Enstatite content, X En ) that can be related to changes in the magmatic system. These changes can lead to either reverse zoning, typically associated with higher temperatures, higher pressures, and higher water fugacity, or normal zoning, typically associated with cooler and shallower environments.

The compositions of unzoned, single zoned and multiple zoned opx along rim-to-rim or core-to-rim traverses allow us to identify compositional plateaus (En mol% plateaus) (Supplementary Table S3). Compositional plateaus have been here defined as regions of the profiles with a minimum length of ∼15 µm (respect to 315-500 µm wide crystals) within which the En content of the crystal shows a constant value, with accepted variability of ±0.5 mol%. The composition of each plateau is then calculated as an average of all values defining the plateau. The occurrence of plateaus can be interpreted as a result of crystal growth in different, stable ME, with residence times in each environment long enough to be recorded by crystal growth, but also short enough that the compositional gradient has not been erased by diffusional equilibration. The repetitive occurrence of plateaus indicates that compositional zoning observed in a given crystal has not been produced by pure fractionation during growth, but by subsequent and fast changes during crystal growth in response to a change from one ME to another, i.e., by changing pressure, temperature, oxygen fugacity or water fugacity, or physically transferring a crystal to a different melt. In this way, assuming that crystal growth occurred sequentially from core to rim, cores represent the first ME recorded by crystals while outer rims represent the last ME. Analysis of compositional plateaus from core to rim allows us to track either movements of crystals and their carrying melts between different ME or modification of storage conditions by external processes (heating, decompression event, mixing, etc.) through the plumbing system until eruption. This procedure has to be repeated on all analyzed crystals in order to define crystal populations, proportions, and the frequently-travelled crystal transfer pathways, before drawing a CSA analysis diagram. This method allows a simple graphical representation that avoids much of the difficulties of combining and interpreting diverse and complicated zoning patterns between crystals of the same sample.

Here we applied this approach to opx to highlight magma transfers and the magnitudes of these transfers at depth between the different ME. This approach is based on the recognition of the diversity of opx composition interpreted as ME changes. Eventually, we can identify the disrupting events at depth responsible for magma remobilization up to eruption. That is also why a ME does not necessarily correspond to a physical lens of magma within the crust; this will be discussed hereafter.

RESULTS

The Diversity of Orthopyroxene Texture and Composition

A total of 1,452 opx were analyzed for the two eruptions (701 for Roseau, 751 for Layou), including their textural features and chemical compositions (Figure 3; Table 1; Supplementary Tables S2,S3).

Textural analysis by high-resolution BSE images shows that, for both eruptions, 85% of crystals of the 315-500 µm fraction are unzoned (Figures 3A-F; Table 1). Opx belonging to larger size classes (500 μm-1 mm) usually show only zonation linked to the melt evolution by fractional crystallization while smaller size classes (250-315 µm) present clear zonations, in the same proportions as the 315-500 µm fraction. Roseau opx show significantly more simple reverse zonation (higher En% in the rim than in the core; 69%-Figures 3B,C,H,I) compared to normal zoning (lower En% in the rim than in the core; 8%-Figure 3A). Some multiple zonations are present (23%) with a prevalence of the "reverse + normal" zonation (Figure 3I; Supplementary Figure S5). Layou opx show more scattered proportions of zonation types than observed in Roseau, despite the similar proportion of zoned crystals (15%): reverse zonations (44%-Figures 3E,F,K,L) are less abundant and normal zoning (28%-Figure 3D) more prevalent in Layou than for Roseau. Of the zoned Layou crystals, 28% show multiple zonation, but with a predominance of "normal + reverse" zonation (Figure 3L), contrary to Roseau.

Core compositions, as well as the unzoned crystals, help us to constrain the equilibrium state of the magma before the disrupting event responsible for magma remobilization; core compositions of each opx are presented in Figure 3. The unzoned opx from Roseau have a range of composition between En 50 and En 54 with a clear mode at En 51-53 (Figure 3A), whereas unzoned opx of Layou encompass a compositional range between En 48 and En 53 , with a moderate mode at En 51-53 (Figure 3D). For zoned opx, core compositions are centered at En 51-53 for 80% of Roseau zoned crystals (Figure 3B) and for about 60% of Layou zoned crystals (Figure 3E), but with a second small peak at En 49-50 (25%). Thus, unzoned opx and the core of zoned opx describe the same dominant composition domain of En 51-53 for both eruptions (Figures 3A,B,D,E). Thus, the unzoned populations are a good proxy for the cores of the zoned populations for both eruptions. The cores have a common origin, and the rims represent dynamic events that happen subsequently, as detailed hereafter.

Orthopyroxene-Hosted Melt Inclusions

Measuring the composition and volatile content in MIs trapped in phenocrysts is one straightforward manner to infer magma storage depth when associated to matrix glass investigation. In addition to the MI study of Balcone-Boissard et al. (2018),a dedicated investigation of a possible correlation between the MI composition and the host opx has been undertaken on 15 MI-opx pairs (Figure 4). Correction of post-entrapment crystallization was considered on the basis of the K D (Fe-Mg) opx-melt (Putirka, 2008) and only represented between 1 and 2 vol% of the MI volume (Supplementary Table S5). For both eruptions, the measured MIs are trapped within opx that encompass the whole compositional range, En 48-57 (Figure 4). Reporting the data together with existing data in the same samples (Balcone-Boissard et al., 2018) shows that there is no systematic variation of MI composition with host-opx core and rim composition (Figure 4).

Fe-Mg Diffusion Modeling in Orthopyroxene

Core-to-rim composition profiles were obtained and modeled for 44 selected zoned opx (21 for Roseau and 24 for Layou). Layou crystals included both normal-and reverse-zoned examples while only reverse-zoned crystals were studied for Roseau (given the dominance of reversely-zoned crystals in this eruption).

Examples of data interpreted as diffusion profiles for Layou and Roseau are shown in Figure 5. As the purpose here is to constrain the last pre-eruptive processes, only the last change in composition -the outermost zone, excluding quench -has been used for timescale determination in each case.

Results obtained on simple zoned crystals (reverse and normal, for Layou and reverse only for Roseau) show a significant peak of Fe-Mg interdiffusion timescales at 3-10 years before eruption (Figures 6A,C; Table 2). For Roseau, the 17 crystals representing the histogram peak and showing timescales under 10 years (Figures 6A,B) allow us to calculate a timescale of ∼4-5 years prior to eruption, with an absolute uncertainty of -3/+ 11 years (Figure 6B). For Layou, the 14 crystals representing the histogram peak for timescales under 10 years (Figures 6C,D) lead to a calculated timescale of ∼3-4 years prior to eruption, with an absolute uncertainty of -2/+9 years (Figure 6D).

DISCUSSION

Pre-Eruptive Magma Dynamics and Identification of the Disrupting Event at Depth

Our starting conception of the architecture of the magmatic system beneath Dominica is a transcrustal magmatic system, comprising a series of magma bodies stored at different depths within the crust that may be interconnected by dykes and influenced by volatile fluxing (Bachmann and Bergantz, 2003;Cashman et al., 2017)o r various regions of a large zoned magma body with different conditions (partially molten rocks, mush, cold storage zone; Edmonds et al., 2019). This creates a variety of ME characterized by differences in pressure, temperature, oxygen fugacity and volatile saturation. For both eruptions, the stable pre-eruptive conditions of the main magma body were constrained by Solaro et al. (2019) to be similar for both eruptions: 850 ± 25 °C, ∼400 MPa (depth of ∼16 km, considering alocalrockdensityof2.45g/cm 3 ), a melt H 2 Ocontentof∼7-8wt% (i.e., slightly H 2 O-undersaturated conditions) and an oxygen fugacity close to NNO +0. 6. Such stable conditions of the main magma body are highlighted by a population of ∼85% of unzoned opx crystals in the 315-500 μm range. Here we discuss for both eruptions what the other ∼15% of zoned opx can tell us about the last magmatic events that triggered the eruption.

These ∼15% of zoned crystals display similar core compositions [in agreement with the stable pre-eruptive conditions determined by Solaro et al. (2019)] but a variety of rim compositions, and are thus evidence of transfers from a common initial environment to other ME, where compositional rims developed. For Roseau opx, the prevalence of Fe-Mg reverse zoning (69% of singly-zoned crystals) suggests few perturbations of the stored magma. For Layou, the complex zonation pattern (∼44% of reversely zoned opx, 28% of normal zoning, and 28% multiple zoning) comparatively testifies a heterogeneous perturbation of the main magma body during the pre-eruptive period. In the following section, these zoning patterns are interpreted in terms of changes in the physical parameters of the storage region, in order to trace the crystal history and the dynamics of the storage system. We then investigate the origin of the last event responsible for magma remobilization prior to eruption. The En content of opx is above all dependent on temperature: a temperature increase causes an increase in En (e.g., Martel et al., 1999;Solaro et al., 2019). Conversely, the En content of opx is less dependent on pressure. Thus, an increase or a decrease in En content may in some cases be interpreted as a heating or a cooling event, respectively, when magmas of different compositions or temperatures are mixing.

Magma Mixing and Reheating

For both eruptions, the relative homogeneity in the compositions of the melt inclusions trapped in opx showing large X En variations from En47 to En57 (Figure 4) suggests a rather homogeneous and buffered silicate melt composition of low temperature (based on Zr content data as evidenced by Balcone-Boissard et al., 2018). In both eruptions, matrix glasses are rhyolitic but slightly less differentiated (silica-and alkali-poorer) than MIs (Balcone-Boissard et al., 2018). Magma mixing processes are able to create X En increase, either by mixing a dacitic magma with a mafic recharge that produces a less differentiated matrix melt or by mixing different parts of the same mush displaying different degrees of melt differentiation. The hypothesis of a mixing with a more mafic magma has been discarded on the bases of i) the melt inclusion compositions, which reflect a rather homogeneous differentiated melt composition trapped within opx, whatever their En content, and also on the basis of ii) the rhyolitic matrix glasses that do not show textural evidences of mixing. The hypothesis of mixing different parts of the same mush is the most realistic, as i) the matrix glasses are slightly less differentiated than MIs, which is incompatible with a pure evolution by decompression-induced crystallization of the MI trapped in the opx and ii) MIs reflect an evolved melt formed at lower temperature than the matrix glass, as evidenced by the Zr saturation in MIs but not in the matrix glasses (Balcone-Boissard et al., 2018).

Investigating the nature of the zoning changes gives insights into the pre-eruptive magma dynamics, as displayed in Figure 7 and detailed hereafter. In agreement with the CSA method, the fluctuations in X En within each single opx crystal have been divided into compositional groups in order to evaluate the magnitude of the heating or cooling event recorded by the crystals. The Roseau eruption shows 2 ME: ME R 1 (En50-54) and ME R 2 (En54-59), with the population dominantly displaying ME R 1 compositions. The dominant zoning is reverse and may be attributed to an increase of temperature. Though ME R 2 -ME R 1 normal zoning is present, it represents a statistically minor component of the population and is typically followed by reverse zoning back into ME R 2. This suggests a dominant reheating process or magma transfer from ME R 2t o ME R 1 with time. This reheating may be linked to the injection of a hotter and slightly less differentiated silicic magma into a rather homogeneous cooler magma reservoir (Figures 8A,B).

For the Layou eruption, 3 ME are described (by order of abundance): ME L 1 (En51-53), ME L 2 (En47-51), ME L 3 (En53-58). The volumetrically dominant body of ME L 1( E n 5 1 -5 4 ) interacts by exchanging with a cooler, slightly smaller body of ME L 2 (En47-51). This could imply alternating heating and cooling of a magma body, or exchange across an interface in a thermally stratified body, as proposed in Figure 8C.T h e r ei sas i g n i ficant interaction between the coldest ME L 2 (En47-51) and the hottest ME L 3 (En54-59), suggesting that the coolest part, likely the walls of the magma reservoir, was interacting with the hottest magma prior to eruption. As for Roseau, the hottest magma may be the injection of magma batches initially stored in an independent reservoir, diagram, the normal zonation linked to the transfer from ME1 L to ME2 L , is related to the convection within the main reservoir of the crystals first stored in the hottest part and then in contact with the coolest part of the reservoir (Figures 8C,D). As both reverse and normal zonation have a timescale of the same order of magnitude (less than 10 years; Table 2), this would imply that the timescales calculated for both processes are too short for accurate discrimination by the method or that both processes are related.

The injection of a new batch of hot magma may be responsible for the magma convection within the intruded reservoir.

The complexity of the pre-eruptive magma dynamics for the Layou eruption confirms the hypothesis of a system heterogeneous in temperature, with parts of the magma stored at lower temperature (allowing quartz crystallization; Solaro et al., 2019). In contrast, the CSA results suggest more direct paths and fewer magma transfers in the case of Roseau magmatic system.

Comparison With Previous Studies

Various studies have already used the Fe-Mg interdiffusion chronometer in opx to retrieve timescales of magma dynamics prior to eruption (Table 3). These include examples of eruptions across a large range of differentiated magma volumes.

Recent petrological studies have shown that the large magma chambers that feed Plinian eruptions at Santorini (Greece; Flaherty et al., 2018) are probably transient features on geological timescales. Magmas are now widely believed to be generated and incubated in bodies of crystal mush, from which they are extracted and accumulated in melt-dominated magma chambers prior to eruption (Bachmann and Huber, 2016).

Conversely to what is generally expected for phenocrysts in magmas, at Santorini, silicic magmas were not the result of a slow cooling over geologically long time periods. By modeling, as here, the diffusive re-equilibration of chemical gradients in zoned opx (feldspars and cpx too), it has been shown that the phenocrysts grew within only a few hundred years, and possibly only a few decades, for the eruption of the Late Bronze Age eruption of Santorini (Flaherty et al., 2018).

The Fe-Mg interdiffusion in opx of the Oruanui eruption (Taupo volcano, New Zealand), involving a magma volume of 530 km 3 , helped to constrain the timing of the decompression event and extraction of melt occurring in the centuries before the eruption (Allan et al., 2013). For the same eruption the whole history of generation, accumulation and extraction of the Oruanui magma body has been discussed by combining diffusion modeling in opx, magnetites and olivines (Allan et al., 2017). The timescales of priming and accumulation of magma prior to the three youngest eruptions (<2.15 ka) of the Taupo volcanic center indicate that a rapid heating and priming of the crystal-mush, from which the final melt-dominant magma body is extracted, occurred in <120 years prior to eruption (Barker et al., 2016). For the Whakamaru eruption (∼340 ka, >1,000 km 3 of magma), the diffusion modeling of Ti in quartz shows that timescales of the final recharge rejuvenating the mush reservoir are in the range of decades before eruption (Matthews et al., 2012).

By studying the geochemical signature of zircon and sanidines, the timescales of magma storage and differentiation of voluminous high-silica rhyolites at Yellowstone caldera (Wyoming) have been investigated (Stelten et al., 2015). Though this is not the same chronometry, the authors argue that eruptible rhyolites are generated by extracting melt and zircons from a long-lived mush of immobile crystal-rich magma. Residence times of the large-volume rhyolites (∼40-70 km 3 ) are ≤1 kyr (conservatively <6 kyr), which suggests that large volumes of rhyolite can be generated rapidly by extracting melt from a crystal mush. As for Santorini, the Yellowstone magmatic system evidences that magma reservoirs spend most of their time in a largely crystalline uneruptible state and that eruptible magma bodies are ephemeral features.

These examples of Santorini, Taupo or Yellowstone, as Dominica, highlight that volcanic systems are characterized by thousands of years of volcanic repose: magma reservoirs spend most of their time in a largely crystalline uneruptible state. Consequently, eruptible magma bodies are ephemeral features, which do not necessarily require multi-decadal timescales for becoming eruptive by magma assembly and mush rejuvenation. Instead, large magma volumes can be mobilized to eruption over relatively short timescales, with no evident correlation between magma volume and timescales. Such highly dynamic volcanic systems bring important implications for ongoing monitoring and hazard mitigation, particularly in areas of high population density (Sparks and Cashman, 2017;Edmonds et al., 2019).

CONCLUSION

The coupled investigation of crystal and melt composition lead us to identify the disrupting event at depth responsible for the magma mobilization, leading to two major explosive eruptions in Dominica. Crystal zonation, interpreted using the revised CSA approach, coupled with diffusion modeling, provides a detailed view of the pre-eruptive architecture and dynamics of the crystal mush prior to eruption. This highlights the diverse complexity of the magmatic system with several magmatic environments and short remobilization timescales. In the last 10 kyrs, several Plinian eruptions in the central part of Dominica have been described and should be investigated to detail the plumbing system dynamics in time and space.

DATA AVAILABILITY STATEMENT

The raw data supporting the conclusions of this article will be made available by the authors, without undue reservation.

Articles, chapitres, entretiens cliniques…

Figure I. 1 : 43 Figure I. 2 :

 1432 Figure I.1: Evolution of the plumbing system views beneath volcanoes. ...…………………….…….43 Figure I.2: Mush replenishment with 6 possible scenarios (modified from Sparks et al. (2019). …….45

Figure I. 5 : 51 Figure I. 6 :

 5516 Figure I.5: Summary of time intervals determined in recent years for a variety of magmatic processes (a) and current applications and relevant timescales for different magmatic compositions (b) (modified from Costa et al. (2021) and Costa et al. (2020), respectively). ……………………………………….51 Figure I.6: Eruptive styles and classification according to the fragmentation index. …………………53 Figure I.7: Different mechanisms of pyroclastic density currents generation (modified from Dufek et al. (2015)).…………………………………………...……………………………………………….56

Figure II. 1 : 73 Figure II. 3 : 76 Figure II. 5 : 79 Figure II. 6 :

 1733765796 Figure II.1: Subduction characteristics. …………………………………………………………….…71 Figure II.2: Subduction zone of the Lesser Antilles. ………………………………………………….73 Figure II.3: Tectonic model for the eastern Caribbean region (modified from Allen et al. (2019)). ….74 Figure II.4: Cross-section along the Lesser Antilles arc showing the crustal structure based on P-wave velocities vP (yellow: upper crustal, green: middle crust, blue: lower crust) (modified from Melekhova et al. (2019)) (a) and the main magma ponding zones beneath the islands from St. Vincent to Montserrat (b) (modified from Balcone-Boissard et al. (in revision)). …….…………………………………..….76 Figure II.5: Maps of Dominica …………………………………….………………………………….79 Figure II.6: Photographs of the studied volcanoes. ……………………………………………………80

Figure I. 2 :

 2 Figure I.2: Mush replenishment with 6 possible scenarios (modified from Sparks et al. (2019)). In (a) and (b): cooling of the emplaced layer leads to the formation of crystals that settle out (as an "adcumulate" layer), density of the remaining magma decreases and ascends, as it becomes buoyant, as a dyke (a) or a plume (b). In (c), crystals are in suspension in the magma but volatile exsolution leads to an instability in the magma, forming a plume. d) Magma recharge can lead to a remobilized mush, that becomes unstable and ascends in the crust or (e) it can mix with the underlying replenished magma, to form a stable layer or unstable buoyant mixed magma (f).

  (Figure I.3).

a)

  The wide range of closure temperatures in minerals can constrain the temperature-time relations of magmatic systems from a few seconds to millions of years, which are the temperatures at which the mineral becomes closed to exchange of the daughter isotope with the environment via volume diffusion (Costa, 2021); b) Time intervals obtained for different elements and minerals for a given temperature (T = 1200°C and 825°C for minerals from basaltic or rhyolitic systems respectively), pressure and oxygen fugacity, corresponding to the quartz-fayalite-magnetite (QFM) buffer and diffusion lengths (Costa et al., 2020). For the orthopyroxenes, the interdiffusion coefficient used is Dohmen et al. (2016).

  or modelling multicomponents diffusion in apatites (Li the positions of the symbols represent the average values of these times recorded by the system. The times corresponding to magma injections/mixing for basaltic volcanoes generally range from a few weeks to years, whereas for volcanoes emitting more advanced magmas the times are in the order of decades or hundreds of years (Costa et al., 2020). Timescales are shown for: Laki (Hartley et al., 2016); Novarupta (Singer et al., 2016); El Hierro (Longpre et al.

Figure I. 6 :

 6 Figure I.6: Eruptive styles and classification according to the fragmentation index. a-d) Hawaiian to Plinian eruptive styles eruptions; a) Hawaiian eruption in September-October 2022 of Piton de la Fournaise (La Réunion), with lava fountaining (< 300 m) along an eruptive fissure in the Enclos and lava flows (picture taken on 19 September 2022 by PGHM and OPVF-IPGP) (Peltier, 2022); b) Strombolian activity at Yasur volcano (Vanuatu), by night, showing an ash-poor explosion with red-hot lapilli and bombs ejected (Taddeucci et al., 2015); c) Vulcanian eruption of Soufriere Hills volcano (Montserrat, 1997), with a fountain collapse generating pyroclastic density currents, with a buoyant plume that rose to about 10 km (photograph from the North-West by P. Cole)

Figure I. 7 :

 7 Figure I.7: Different mechanisms of pyroclastic density currents generation (modified from Dufek et al. (2015)). They are generated by gravitational collapse of material ejected in the atmosphere; when the bulk density exceeds that of the ambient atmosphere and are generated by several eruptive styles: (a) Directed blasts, when an edifice collapse rapidly decompresses a magmatic body, generating a lateral explosion (Mount St. Helens, USA, in 1980) (Christiansen and Peterson, 1981); (b) Dome collapse, when a dome's interior is decompressed, producing a directed explosion (Soufriere Hills, Montserrat, in 1997) (Druitt et al., 2002); (c) Boiling-over eruptions or low column collapse, occurring when vesiculated magma fountains go over the crater rim without forming a convective plume (Tungurahua, Ecuador, in 2006) (Hall et al., 2013) and (d) Column or fountain collapse, when the fragmented mixture fails to entrain sufficient atmospheric gases to become buoyant, collapsing on itself and forming pyroclastic density currents.

  (Figure I.9a). eruption) and repose (time since the previous eruption) (b) (modified from Ruth et al. (2018) and Rassmussen et al. (2018), respectively). a) Combination of geochemical, seismic and timescales prior to the 2008 eruption of Llaima (Chile) using olivine diffusion chronometry. Magma intrusion occurred 4 years before the eruption and then magma migration happened, reaching 4 km depth 6 months before the eruption, which can be linked to reported seismicity and ash emissions (Ruth et al., 2018). b) Hot spot and arc volcanoes are circled in purple or blue, respectively. Eruptions that occurred at "open-systems" volcanoes (no pre-eruptive deformation recorded) show a longer run-up time (circled symbols), compared to other volcanoes. ME-1, 2 and 3 are mixing events that preceded the 1999 eruption of Shishaldin volcano; for volcano abbreviations significations, see Rasmussen et al. (2018).

Figure II. 1 :

 1 Figure II.1: Subduction characteristics. a) Schematic section through the upper 140 km of a subduction zone, with the interactions between the crustal and upper mantle components (Stern, 2002). The mantle wedge is located in the part of the mantle beneath the overriding plate and between the trench and the most distal part of the arc where subduction-related igneous or fluid activity is found (modified from Stern (2002)). b) Diagram showing the geochemical characteristics in trace elements of OIB (Sun and McDonough, 1989), N-MORB (Normal Midocean ridge basalts) (Sun and McDonough, 1989) and arc lavas (Monzier et al., 1997) (modified from Hoffer (2008)). The compositions are normalized over the primitive mantle of Sun and McDonough (1989).

Figure II. 2 :

 2 Figure II.2: Subduction zone of the Lesser Antilles. a) Three-dimensional diagram showing the Northern and Southern American plates in curved and oblique subduction under the Caribbean plate (modified from the GARANTI campaign © Géosciences Montpellier and from Favier (2019)). b) Schematic map of the Lesser Antilles showing the three volcanic segments (north, central, and south). The inner and outer arc are also shown, with active volcanic islands and older volcanic islands mostly covered by coral reef platforms, respectively (Wadge and Shepherd, 1984) and the exposure of island populations (circles based on the number of people living within 15 km of a volcano summit) (modified from Leone et al. (2019)). Volcanoes in the central part generally emit differentiated magmas with high explosive potential (Leone et al., 2019; Macdonald et al., 2000). Approximately one million people would be threatened by the effect of direct volcanic activity in the Lesser Antilles (such as pyroclastic density currents, lava domes, directed blasts…) and related phenomena would threaten 2.5 million people (lahars, tsunamis, volcanic gases…) (Lindsay et al., 2005a). Morne Watt is a satellite volcano of Morne Trois-Piton Micotrin.

Figure II. 3 :

 3 Figure II.3: Tectonic model for the eastern Caribbean region (modified from Allen et al. (2019)). a) Around 60 Ma, the regional arc magmatism is associated with the late stages of the Great Arc of the Caribbean (GAC), evidence for this arc is found in the rock record of several island groups, including the Dutch, Venezuelan, and Greater Antilles, as well as Aves Ridge. CLIP-Caribbean large igneous province. b) Slab rollback led to abandonment of the Aves Ridge and eastward migration of arc volcanism at 40 Ma. New Outer arc includes the northern islands of the Lesser Antilles arc ("Limestone Caribbees" on (b)) and the backarc spreading is only present in the south, with a north-south-oriented spreading axis. c) 0 Ma: shallowing of the slab leads to the abandonment of the Outer arc and establishment of the modern Lesser Antilles arc to the west (Allen et al., 2019).

  II.4b) (Balcone-Boissard et al., in revision).

  (Figure II.8b).

Figure II. 8 :

 8 Figure II.8: Location of Kamchatka. a) Pacific Ring of Fire (modified from Britannica. The Editors of Encyclopaedia (2011)). Oceanic trenches are in blue and the volcanic islands are parallel to the trenches. b) Main geographical and tectonic units of the Kuriles-Kamchatka-Aleutian-Alaska subduction, with a topographic/bathymetry base map (modified from Koulakov et al. (2011)). Tectonic-plate boundaries are in orange and violet ones are drawn by Koulakov et al. (2011). WB: Western Basin; KOB: Kamchatka-Okhotsk block; OB, Okhotsk block. The Kamchatka Peninsula lies above the northwestern margin of the Pacific Plate dipping at a rate of ~8 cm/yr (DeMets, 1992) beneath accreted Mesozoic and Tertiary volcanic terraces, forming the Kamchatka volcanic arc from 51°N to 60°N (1250 km long) (Figure II.8b; Figure II.9a)(Gorbatov et al., 1997;[START_REF] Kayzar | Deciphering petrogenic processes using Pb isotope ratios from time-series samples at Bezymianny and Klyuchevskoy volcanoes, Central Kamchatka Depression[END_REF][START_REF] Konstantinovskaia | Arc-continent collision and subduction reversal in the Cenozoic evolution of the Northwest Pacific: an example from Kamchatka (NE Russia)[END_REF][START_REF] Koulakov | Slab shape in subduction zones beneath the Kurile-Kamchatka and Aleutian arcs based on regional tomography results[END_REF] Minster and Jordan, 1978;[START_REF] Watson | Tectonic evolution of Kamchatka and the sea of Okhotsk and implications for the Pacific Basin[END_REF]. The Pacific plate entering the Kamchatka trench has an inferred age of ~95 Ma at the southern tip of Kamchatka(Portnyagin et al., 2007a;[START_REF] Renkin | Depth and age in the north Pacific[END_REF]. For some authors, Kamchatka and eastern Siberia constitute the western end of the Northern American Plate

  (Figure II.9b). Numerous approaches explained one of each of these features but the model that explains the entire dataset is the one of Lander and Shapiro (2007) described below.

Figure II. 9 :

 9 Figure II.9: Simplified map of Kamchatka and its volcanoes (Kamchatka Volcanic Eruption Response Team (KVERT), 2002) (a) and 3D visualizations (b-c). Numbers correspond to the three peninsulas in the eastern coast of the Kamchatka segment (1 -Shipunski, 2-Kronotski and 3-Kamchatkaski Mys). b) 3D visualizations of the Kamchatka subduction zone seen from the north (modified from Portnyagin et al. (2008)). Dashed lines and colors (blue to red) show the increase in depth of the plunging panel (in km). The light-colored arrow shows the direction of Pacific plate motion. Tolbachik (T), Klyuchevskoy (Kl), Zarechny (Z), Kharchinsky (Kh), Shiveluch (Sh) and Shishelsky (Shs) volcanoes are shown. c) Subduction zone with the plunging plate in blue (modified from Levin et al. (2002)). Active (pink) and inactive (green) volcanic arcs are shown.

Figure

  Figure II.10: Post-Eocene Kamchatka subduction zone history and geodynamic reconstructions (a1-a4) from 38 to 2 Ma (modified from Lander and Shapiro (2007)). (1. Volcanic belts, 2. Trenches, 3. Transform faults, 4. "Grechishkin thrust" suture, 5. Pacific plate velocity vector relative to Eurasia). The ages of the collision are estimated from data on sediment ages, paleomagnetic estimates and maximum depths of earthquakes in different parts of the modern Kamchatka seismic zone by deducing the amount of time since the beginning of subduction from the dip extension of the slab (the Kamchatka segment is divided in three smaller zones, as the depth of the seismic zone is not constant) (Lander and Shapiro, 2007).

  (Figure II.13a). At the same time, the tomographic model by Koulakov et al. (2020) revealed an interconnected feeding system of the CKD volcanoes (Klyuhevskoy group and Shiveluch Yerov); c) Picture of the volcano after the 1956 flank collapse and its horseshoe-shaped structure, in August 1956 (photo by G. S. Gorshkov); d) Bezymianny in 2019, with lava domes filling the horseshoe shaped structure (photo by Yu. V. Demyanchuk); e) Sketch map of the deposits of the 1956 climactic phase showing the boundaries of the debris avalanche and the laterally directed blast deposits (Belousov, 1996).

  (Figure II.16g). This style continued until late 2012 (Figure II.16h) and the volcano remained dormant between September 2012 and December 2016. This short period of dormancy ended with the onset of an effusive eruption with lava dome growth (December 2016) and two strong explosive eruptions in March and June 2017 (Mania et al., 2019) (Figure II.16i). Thereafter, Bezymianny returned to short explosive eruptions approximately once a year, the last of which was in August 2022 and generated pyroclastic density currents.

  (Figure II.14b; Figure II.15d).

Figure II. 16 :

 16 Figure II.16: Bezymianny volcano general eruptive timeline from 1956 to present-day. After a long dormancy, reactivation of Bezymianny volcano on 22 October 1955 with Vulcanian-type ash eruptions from 22 October until the end of November, with (a) Picture of Bezymianny on 30 October 1955 (preceded by a 23-days seismic swarm); b) Fumarolic activity on 23 January 1956 during lava dome growth and (c) Flank collapse at Bezymianny that generated a lateral directed blast and Plinian activity on 30 March 1956 (view from Kozyresk village, 45 km away from the volcano). (d-e) Before and after the eruption, (d) in July 1949, with the black dashed line showing the outline of the volcano in February 1956 and red dots, crater in October 1955. (e) Bezymianny after the explosion in March 1956, with the black line showing the outline of the volcano before the eruption. f) Consequences of the 1956 explosion: the bushes and forest trees were broken and fell away from the volcano over an area of ~500 km 2 at the eastern foot of Bezymianny at a distance of 25 km from the volcano (Belousov and Belousova, 1998), example in the valley of the

  (Figure II.13).

  [START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF] also indicated the presence of two distinct reservoirs from amphibole thermobarometry, with deeper depths than what was inferred by[START_REF] Shcherbakov | Phase equilibria constraints on preeruptive magma storage conditions for the 1956 eruption of Bezymianny Volcano, Kamchatka, Russia[END_REF]: between 8-11 km (200-300 MPa, 900 ± 50 °C) and[19][20][21][22][23][24][25][26] 975 ± 25 °C) (Figure II.17).

  (Figure II.17a-b). Magmas from the shallowest reservoir erupted first (200-300 MPa), with increasing proportions of deeper reservoirs with time (up to 2010) (Figure II.17a-b). The 1956 climactic phase was likely to have been triggered by a mafic injection at depth because the composition of the lower reservoir was likely modified by the injection of a melt from the high-pressure reservoir that became more mafic (indicating an injection with a third, even more mafic liquid (end-member mixing analysis)) (Figure II.17b).

Figure II. 17 :

 17 Figure II.17: Cartoon model of physical processes prior to 1956, with three mixing components of Bezymianny evolving apart from each other via crystal fractionation (modified from Turner et al. (2013)). a) Before 1956, two magma reservoirs are present and isolated (end-members A and B), evolving via crystal fractionation from a common source. b) In 1956, a magma injection (magma C) into the lower chamber triggers the eruption, magmas contained in the shallower

Figure II. 18 :

 18 Figure II.18: Seismic stations and Klyuchevskoy group. a) Location of the first seismic station installed near Bezymianny (Apakhonchich, 16 km northeast of Bezymianny, blue square) (Bogoyavlenskaya et al., 1991; Girina, 2013); ASTER satellite image on 8 February 2005 (NASA/GSFC/METI/Japan Space Systems, and U.S./Japan ASTER Science Team) (modified from Girina (2013)). b) Bezymianny Volcano in the Klyuchevskoy volcanic group, Kamchatka and surrounding volcanoes (red triangles) (picture taken in August 2019).

Figure II. 19 :

 19 Figure II.19: Map of the deposits of the 2010-2013 eruption of Kizimen (image: NASA Operational Land Imager acquired on 2 September 2013) (modified from Auer et al. (2018)). Dotted line delineates the areas of the block and ash pyroclastic density current deposits (eastern fan (I), north-eastern fan (II), southern fan (III)); dashed line delineates the thick lava flow (maximum thickness of lava flow front of ~232 m (Dvigalo et al., 2013)). Inset shows location of Kizimen on the Kamchatka peninsula.

  (Figure II.19). Signs of reactivation of the volcano were observed on 16 October 2010 with the formation of two fumarole fields at the top of the snow-covered Successful prediction of a strong explosive eruption of Kizimen volcano (VEI = 3) was made, despite the fact that no historical magmatic eruption had been acquired on this volcano, thanks to the previous experience on Kamchatka volcanoes and literature data on similar volcanoes of the world[START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF] (Figure II.21c-f).

Figure II. 21 :

 21 Figure II.21: Network of the closest seismic stations of Kizimen volcano (modified from Sharikova and Firstov (2021)) (a-b) and seismic activity of Kizimen volcano from 01.01.1962 to 01.05.2011 (modified from Senyukov et al. (2011)) (c-f). a) Triangles showing seismic stations ("TUM" -Tumrok, installed in July 2003, "KZV" -Kizimen station, considered in data processing from 12 November 2009 and "TUMD", which started recording in July 2011 (Senyukov et al., 2011; Shakirova and Firstov, 2022)); b) Distribution of the volcano-tectonic earthquakes, depending on

. 22 )

 22 . Then, the hybridization of the base of the dacitic magma occurs with the basalt, forming an andesitic magma of hybrid composition. Then, after a long period of dormancy, a phase of seismic activity started in April 2009 and continued for one and a half years before the eruption[START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF] (Figure II.21c-e; Figure II.22).

  (Figure II.21a). In July 2003, a new radiotelemetry seismic station was installed 20 km to the north-west of the volcano, increasing the seismic record on Kizimen (TUM-Tumrok) (Senyukov et al., 2011) (location shown on Figure II.21a). In 2009, another station was installed on the southeastern slope of the volcano, 2.5 km away from the summit to better assess the volcanic hazard and increase the representativity of the earthquakes (KZV; location shown on Figure II.21a), as a strong explosive eruption of Mount St. Helens 1980 type could occur on this volcano (Melekestsev et al., 1995; Senyukov et al., 2011) and it was considered in data processing from 12 November 2009 (KZV-Kizimen station) (location shown on Figure II.21a). It improved the representative classes of the earthquakes, but also increased the accuracy of determining the origin positions of earthquakes (Senyukov et al., 2011) (Figure II.21). In July 2011, a new seismic station located 8 km away from the volcano (TUMD; Figure II.21a) started recording and improved the representativity of local magnitude

Finally, 6 1 :

 61 samples from the current active dome as well as 2 samples from a March 2019 dome collapse were also collected (Figure III.Localization and nature of the studied samples of Bezymianny.

Figure III. 2 :

 2 Figure III.2: Sketch maps of the eruptive deposits and samples localization. a) Bezymianny, with the boundary of the debris avalanche and the laterally directed blast deposits from 30 March 1956 climactic phase of the eruption, following Belousov (1996). Image credit: Google Earth V 7.3.4.8248 (July, 2013). Bezymianny, Kamchatka, 55° 56.880'N, 160° 40.282'E, eye altitude 17.57 km. Image Landsat/Copernicus. Maxar Technologies 2022. [April 12, 2021]. b) Kizimen, with the 2010-2013 eruptive deposits. Dotted line delineates the area covered by the block and ash pyroclastic density current deposits and the dashed line, the limit of the lava flow (following Auer at al. (2018)). Image credit: Google Earth V 7.3.4.8248 (April, 2006). Kizimen, Kamchatka, 55°

  . The different stages of sample preparation are illustrated in Figure III.7.

Figure III. 7 :

 7 Figure III.7: Preparation of samples from the rock sample to the crystal mounts of different types. Selection of the samples of each eruption; some part will be used for thin sections (a), whole rock analysis (b) or textural and chemical analysis (c-g). c) Cleaning and crushing of the samples; d) sieving; e) Sorting by hand under binocular magnifier of the crystals in the various fractions (orthopyroxenes surrounded in brown, plagioclases surrounded in grey, amphiboles in black and magnetites surrounded in black in e1-4); f) Crystal mounting, impregnation and polishing, with an example of orthopyroxenes mount (overview realized with Scanning Electron Microscope (SEM)) after mounting of these crystals along the c-axis, in the north-south direction for the modelling of the intracrystalline interdiffusion (see part II. 3) and (g) polishing in order to carry out analyses with the SEM and the electron microprobe.

Figure III. 11 :

 11 Figure III.11: Examples of defects in crystals (modified from Costa and Morgan (2010)). a) Schottky defect are gaps compensated by charges on different sites of the crystal system, which can migrate independently but retain the overall neutrality of the charge of the crystal. b) Frenkel defect, created by migration of an ion from a crystal site either to an adjacent site, or by maintaining this ion in an interstitial position, generating a vacancy at the original position. In the case of ions with small ionic radii and small charge, this mechanism is the easiest (Li ion for example).

  (Figure III.14).

Figure III. 13 :

 13 Figure III.13: Initial zoning profiles used in the diffusion time modelling (dotted lines) and their evolution after a given diffusion time (solid lines) (modified from Costa et al. (2008)). Four types of initial conditions can be identified: a) Smooth zoning that can be created following magmatic differentiation; b) Abrupt zoning created following changes in the storage conditions of the magmas, faster than crystal growth (for example, magma mixing event); c) Homogeneous zoning where diffusion happens after a later change of concentrations at the boundary; d) Oscillatory zoning, for example in the case of rapid changes of storage conditions in the magmatic reservoir.

Figure III. 14 :

 14 Figure III.14: Illustration of the use of boundary conditions in closed (a-b) or open (c-d) systems (modified from Costa et al. (2008)). a) The crystal is isolated from its environment (no flux); c) a flux can be present at the edge of the crystal. The evolution of the profiles under these two types

  MH (magnetite-hematite): 4 Fe3O4 + O2 = 6 Fe2O3 -RI (rutile-ilmenite): 2 Fe2O3 + 4 TiO2 = 4 FeTi O3 + O2 -NNO (nickel-nickel oxyde): 2 Ni + O2 = 2 NiO -FMQ (fayalite-magnetite-quartz): 3 Fe2SiO4 + O2 = 2 Fe3O4 + 3 SiO2 -WM (wustite-magnetite): 3 Fe1-xO + O2 ~ Fe3O4, with x ~1 -IW (iron-wustite): 2(1-x) Fe + O2 = 2 Fe1-xO -QIF (quartz-fer-fayalite): 2 Fe + SiO2 + O2 = Fe2SiO4

  following order: D (𝐹𝑒-𝑀𝑔)𝑐 >> D (𝐹𝑒-𝑀𝑔)𝑏 >> D (𝐹𝑒-𝑀𝑔)𝑎 , where D (𝐹𝑒-𝑀𝑔)𝑥 is the diffusion coefficient of Fe-Mg along a specific axis x (Ganguly and Tazzoli, 1994) (Figure III.17b). So, Fe-Mg interdiffusion should be fastest along the c-axis and slowest along the a-axis and the overall kinetics of Fe 2+ -Mg fractionation (order-disorder processes) between the M1 and M2 sites should be dominated by exchange along the c-and b-directions.

  2 mol %). The profiles are acquired perpendicular to the caxis, parallel to the b-axis of the crystals and far from the crystal ends(Figure III.19). Indeed, profiles acquired along the c-axis of the crystals can reveal anomalies, being vulnerable to recrystallization(Allan et al., 2013; Solaro-Müller, 2017). Profiles in grayscale intensities are then obtained from the BSE images with ImageJ software (https:// imagej. nih. gov/ ij/; version1.52a) and averaged over a certain width to reduce the noise(20-50 µm). b) Initial and boundary conditions: -Initial condition: abrupt zonation (Figure III.19) -Boundary conditions: open, which implies an exchange of chemical elements with the surrounding magma and the variation of the composition of the crystal boundary. c) Model:

Figure III. 22 :

 22 Figure III.22: Plumbing system of Etna (modified from Kahl et al. (2011)). a-b) Simplified systems analysis of the zoning patterns in olivines from Etna eruptions in 1991-1992. Each magmatic environment is represented by a box and arrows between them indicate the passage of the recorded environments, from the core of the crystal to the rim (Kahl et al., 2017, 2015, 2013, 2011). The density of the connection lines can be used to infer the dominant passageways of magma transport. c) Magma dynamics coherent with a) and b) with the different timescales (Δt1, Δt2 and Δt3) that can be extracted from diffusion modelling which are linked to duration of residence of melts in different segments of the plumbing system.

  Figure III.23a-c). A second melting process was performed to obtain completely homogeneous glasses, in a larger crucible, which was used afterwards for the experiments (77 mm high, 37 mm wide) (Figure III.23d).

  . The experimental setup consists of a Pt80Rh20 crucible (outer cylinder) filled with glass of one specific composition and a core of the other glass composition to be mixed, as well as a spindle (inner cylinder) eccentric to the center of the outer cylinder that will allow internal rotation during the experiment (Figure III.23f). The geometry of COMMA is defined by two parameters: (1) the ratio of the radii between the two cylinders (𝑟 =

Figure III. 23 :

 23 Figure III.23: Crucibles containing the magmatic melts. a-c) First melting cycle of the first endmember in the furnace (at 1500 °C) and sudden cooling in the air to produce glass. d-f) After the second melting, crucibles and the two end-members glasses produced. d) A core of the melt B is put in the crucible with melt A. e) Focus on the experimental crucible with the melt A composition where a glassy cylinder of melt B composition is inserted. f) The off-center position of the spindle is also shown. The main experimental parameters are Rout (18.5 mm) radius of the outer cylinder with the dacitic melt, Rin (6.2 mm) radius of the inner cylinder sheated with the spindle, Rm radius of the basaltic melt (5.5 mm), Ωout (0.11 rpm) angular velocity of the outer cylinder, Ωin (0.66 rpm) angular velocity of the inner cylinder (Morgavi et al., 2015; Rossi et al., 2017). g) COMMA apparatus; h-i) Pictures of the crucible after the end of the experiment.

  of the most magmatically active islands of the Lesser Antilles through its four active volcanoes, is likely host under its central part, below Morne Trois Pitons-Micotrin, to a wellestablished transcrustal mush system. Pre-eruptive spatiotemporal magma dynamics are examined for five, explosive, pumiceous eruptions of this volcano in the last 24 kyrs through a combined Crystal System Analysis and intracrystalline Fe-Mg interdiffusion timescales modelling approaches. Before all eruptions, two magmatic environments of close compositions have interacted. These interactions began ~ 10-30 years prior to the four smaller of these eruptions, with more sustained mixing in the last decade, accelerated in the last 2 years. This contrasts with the largest pumiceous eruption, involving deeper magmas, with magma interaction starting over roughly a century but with various patterns. This suggests a possibility that increasing reactivation signals could be registered at the surface some years before future eruptions, having significant implications for volcanic risk mitigation.

Figure 1 .

 1 Figure 1. Dominica eruptions. (a) The Lesser Antilles arc60-62 results from the subduction of the northern and southern Atlantic plates beneath the Caribbean plate at a rate of ~ 2 cm/yr63 . Black: volcanic islands, grey: coral platforms; the map has been modified from a previous study39 . Bathymetry is in light grey60 . Inset: Shuttle Radar Topography Mission (SRTM) topography of Dominica with four actives volcanic centers (Morne aux Diables, Morne Diablotins, Morne Trois Pitons-Micotrin and Morne Plat Pays, from north to south) and the largest magmatic production of the Lesser Antilles arc (~ 40 km 3 for the last 100 kyrs) (Courtesy of Dr. Ian C.F. Stewart). (b) Map of southern Dominica and location of the studied deposits. From the Morne Trois Pitons-Micotrin volcano, in purple: Grand Fond (GF) large explosive eruption deposits in the valley, outcrop at (15°21′57″N, 61°15′16″W, marked by a star on the map) and Goodwill Entrance (GE), PPR1, 2 and 3, in the Roseau Valley (Goodwill quarry: 15°18′32″N, 61°23′03″W, star on the map), in grey: Roseau ignimbritic deposits 34 . The maps of (a,b) have been done with CorelDraw Graphics Suite 2020 (https:// www. corel draw. com/ fr/). (c) Chronology of the pumiceous eruptions from Morne Trois Pitons-Micotrin for 70 kyrs. Rectangles: filled ones: eruptive deposits from Morne Trois Pitons-Micotrin and empty ones: eruptive deposits from Morne Diablotins, their height depends on the magma volume (km 3 )34 . Colors: purple for Grand Fond, blue for Goodwill, yellow for PPR1, red for PPR2, green for PPR3 and grey for the unstudied eruptions. The large explosive eruptions involve 4-5 km 3 of magma Dense Rock Equivalent (DRE); the voluminous pyroclastic density currents, which are several tens of meters thick in the main valleys, are more or less welded. The small explosive eruptions involve magma volumes of at least one order of magnitude lower than the large explosive ones34 .

  Figure 1. Dominica eruptions. (a) The Lesser Antilles arc60-62 results from the subduction of the northern and southern Atlantic plates beneath the Caribbean plate at a rate of ~ 2 cm/yr63 . Black: volcanic islands, grey: coral platforms; the map has been modified from a previous study39 . Bathymetry is in light grey60 . Inset: Shuttle Radar Topography Mission (SRTM) topography of Dominica with four actives volcanic centers (Morne aux Diables, Morne Diablotins, Morne Trois Pitons-Micotrin and Morne Plat Pays, from north to south) and the largest magmatic production of the Lesser Antilles arc (~ 40 km 3 for the last 100 kyrs) (Courtesy of Dr. Ian C.F. Stewart). (b) Map of southern Dominica and location of the studied deposits. From the Morne Trois Pitons-Micotrin volcano, in purple: Grand Fond (GF) large explosive eruption deposits in the valley, outcrop at (15°21′57″N, 61°15′16″W, marked by a star on the map) and Goodwill Entrance (GE), PPR1, 2 and 3, in the Roseau Valley (Goodwill quarry: 15°18′32″N, 61°23′03″W, star on the map), in grey: Roseau ignimbritic deposits 34 . The maps of (a,b) have been done with CorelDraw Graphics Suite 2020 (https:// www. corel draw. com/ fr/). (c) Chronology of the pumiceous eruptions from Morne Trois Pitons-Micotrin for 70 kyrs. Rectangles: filled ones: eruptive deposits from Morne Trois Pitons-Micotrin and empty ones: eruptive deposits from Morne Diablotins, their height depends on the magma volume (km 3 )34 . Colors: purple for Grand Fond, blue for Goodwill, yellow for PPR1, red for PPR2, green for PPR3 and grey for the unstudied eruptions. The large explosive eruptions involve 4-5 km 3 of magma Dense Rock Equivalent (DRE); the voluminous pyroclastic density currents, which are several tens of meters thick in the main valleys, are more or less welded. The small explosive eruptions involve magma volumes of at least one order of magnitude lower than the large explosive ones34 .

Figure 2 .

 2 Figure 2. The different zonations identified in the opx cargo: images and proportions. SEM images (Sorbonne Université, Paris). (a) Unzoned crystal. Two types of single-zoned opx (SZ) have been identified: normal-zoned (NZ, b), with a Mg-rich core (darker zone on a grayscale image) and Fe-rich rims (clearer in grayscale), or reverse-zoned (RZ, c), with inversely a Fe-rich core and Mg-rich rims. (d) Multiple-zoned opx (MZ), either reverse + normal (R + N) or normal + reverse (N + R) or with a third rim (N + R + N or R + N + R, Supplementary FigureS3). The red line shows the location of the electron microprobe profiles (EPMA) and the core, inner ("i") and outer rim ("o") zones are annoted on the multiple-zoned opx(d). From (e) to (i) (drawn with Microsoft Office suite 2019 Version 1808; https:// www. micro soft. com/ fr-fr/ micro soft-365): proportions of the zonations identified in the opx of Grand Fond in the 355-500 µm fraction (e) and in the opx of the small explosive eruptions of the three fractions studied all together (355-500, 250-355 and 125-250 µm fraction, f to i). 4% of the multiple zoned crystals studied had a third rim (less than 5 crystals per eruption). However, this third rim, though existing in the crystal, does not form a plateau in the compositional profiles (< 10 µm), so they will not be considered in the following results (Supplementary FigureS3). These last rims could correspond to very shorts events prior to the eruptions.

Figure 4 .

 4 Figure 4. Frequency histograms of the En content of zoned opx of the five studied eruptions. The two populations of opx identified are the same as in Fig. 3. Population 1: cores ≥ En 53 for Grand Fond, En 54 for Goodwill and En 52 for PPR1, 2 and 3 and population 2: cores < En 53 for Grand Fond, En 54 for Goodwill and En 52 for PPR1, 2 and 3. The color for each eruption is the same as in Fig. 3. (a1) to (a5) single-zoned (SZ) opx of population 1; (b1) to (b5) multiple-zoned (MZ) opx of population 1; (c1) to (c5) SZ of population 2; (d1) to (d5) MZ of population 2. The arrows underline the major path from the core to the rim(s) recorded by the SZ opx or, for MZ, from the core to the inner rim and dashed line: inner rim to the outer one of the MZ opx. The magmatic environments (ME) discussed afterwards are specified here for each eruption. n: number of opx analysed.

Figure 5 .

 5 Figure 5. Frequency histograms of Fe-Mg interdiffusion timescales modelled in opx at 850 ± 25 °C for Grand Fond (a,b) and 890 ± 9 °C for core-rim and inner-outer rim boundaries of single-zoned (SZ) and multiple-zoned (MZ) opx of the two opx populations as inFigs. 3 and 4, of Goodwill (c,d), PPR1 (e,f), PPR2 (g,h) and PPR3 (i,j). The type of zonation is specified (reverse-zoned or normal-zoned). The timescales are presented as times before the eruptions and can be found in Supplementary TableS3. The colors are the same as in the previous figures. n is the number of profiles modelled (in SZ and MZ opx). The number "n" can be different from the total number of timescales modelled as timescales in MZ opx are sometimes modelled on both the inner and outer rims (e.g.: two timescales modelled in one profile). Individual times and their absolute uncertainties based on the propagation of a temperature uncertainty of ± 25 °C12,38,44 or 9 °C39 and uncertainties associated to SEM images are presented in Supplementary FigureS4.

Figure 6 .

 6 Figure 6. Crystal System Analysis approach of the zonations recorded by the two populations of opx of the five eruptions studied [Grand Fond (a), Goodwill (b), PPR1 (c), PPR2 (d) and PPR3 (e)]. Each box represents a magmatic environment (ME): ME1 GF = En 53-56, ME2 GF = En 49-53 ; ME1 GE = En 54-58, ME2 GE = En 50-54; ME1 PPR1 = En 52-59, ME2 PPR1 = En 46-52 ; ME1 PPR2 = En 52-58, ME2 PPR2 = En 48-52 ; ME1 PPR3 = En 52-57, ME2 PPR3 = En 48-52,for Grand Fond, Goodwill, PPR1, 2 and 3, respectively. For each eruption, the size of the boxes is approximately proportional to the proportions of unzoned and cores of zoned opx in each ME, using the volumes of ME calculated in Supplementary TableS4. The grey zones show the proportions of unzoned crystals and the white areas, the proportions of zoned crystals (reset to 100%, Supplementary TableS4). The change of ME recorded from the core to the rim of opx is represented with a black arrow for single-zoned (SZ) crystals and a red arrow for multiple-zoned (MZ) crystals. For PPR3 (e), 13 multiple-zoned crystals record a change from ME1 PPR3 to ME2 PPR3 (13 in red). In the case of the studied magmas, an En increase (reverse-zoned) implies a temperature increase, with, or without, magma mixing with a more basic and hotter magma. Conversely, an En decrease (normal-zoned) may be linked to a pressure drop, decrease in melt H 2 O content or mixing with a colder magma38 . This Figure has been drawn using Adobe Illustrator (Version 24.0.1; https:// www. adobe. com/ fr/ produ cts/ illus trator. html).

Figure 7 .

 7 Figure 7.Simplified view of the storage areas and of the pre-eruptive dynamics of the magmas in the magmatic mush for Morne Trois Pitons-Micotrin. Models based on opx compositions and interdiffusion timescales modelling for the small explosive eruptions and Grand Fond. The different magma environments recorded are represented by magma lenses at different depths in a transcrustal system. This schematization agrees with the vision of the magmatic mush showing magma lenses stored at different depths6 . On top: model of the whole crust with a deep reservoir feeding the large explosive eruptions and a shallower one feeding the small explosive eruptions38,39 . (a,b) initial stages for small explosive eruptions (Goodwill, PPR1-3) and voluminous pumiceous eruption of Grand Fond; c: rank ordering of the interdiffusion timescales modelled in the opx of Goodwill, PPR1-3 (Fig.5; Supplementary TableS3). Each curve in (c) can be divided into three gradients (underlined by black arrows): a shallow one (from 32 to 10 years), a steep one (from ~ 10 to 2-3 years), together with that of a pre-eruptive region (timescales from 0 to 2-years, vertical arrow). A steep gradient can indicate a rapid processing whereas a shallow gradient, a much slower processing.(d,e) evolution of the magmas prior to the small explosive eruptions with vigorous magma mixing in a temperature-zoned reservoir, with ME2 being the wall and ME1 the center of the reservoir, in the different periods of time underlined in (c); (f) rank ordering of the interdiffusion timescales modelled in Grand Fond and Roseau 12 , large explosive eruptions from Morne Trois-Pitons Micotrin volcano compared to the small explosive eruptions (Fig.5; Supplementary TableS3). In (f), for Grand Fond, one gradient can be identified (from 66 to 5 years). For Roseau, two gradients can be identified: a shallow one (from 60 to 7 years) and a steeper one from 7 to 1 year before the eruption. (g,h) evolution of the magmas prior to the large explosive eruption with magma mixing following the period of times shown in (f). T: temperature; black arrow: magma mixing. SZ: single-zoned opx and MZ: multiple-zoned opx. This Figure has been drawn using Microsoft Office suite 2019 Version 1808 (https:// www. micro soft. com/ fr-fr/ micro soft-365).

  Figure 7.Simplified view of the storage areas and of the pre-eruptive dynamics of the magmas in the magmatic mush for Morne Trois Pitons-Micotrin. Models based on opx compositions and interdiffusion timescales modelling for the small explosive eruptions and Grand Fond. The different magma environments recorded are represented by magma lenses at different depths in a transcrustal system. This schematization agrees with the vision of the magmatic mush showing magma lenses stored at different depths6 . On top: model of the whole crust with a deep reservoir feeding the large explosive eruptions and a shallower one feeding the small explosive eruptions38,39 . (a,b) initial stages for small explosive eruptions (Goodwill, PPR1-3) and voluminous pumiceous eruption of Grand Fond; c: rank ordering of the interdiffusion timescales modelled in the opx of Goodwill, PPR1-3 (Fig.5; Supplementary TableS3). Each curve in (c) can be divided into three gradients (underlined by black arrows): a shallow one (from 32 to 10 years), a steep one (from ~ 10 to 2-3 years), together with that of a pre-eruptive region (timescales from 0 to 2-years, vertical arrow). A steep gradient can indicate a rapid processing whereas a shallow gradient, a much slower processing.(d,e) evolution of the magmas prior to the small explosive eruptions with vigorous magma mixing in a temperature-zoned reservoir, with ME2 being the wall and ME1 the center of the reservoir, in the different periods of time underlined in (c); (f) rank ordering of the interdiffusion timescales modelled in Grand Fond and Roseau 12 , large explosive eruptions from Morne Trois-Pitons Micotrin volcano compared to the small explosive eruptions (Fig.5; Supplementary TableS3). In (f), for Grand Fond, one gradient can be identified (from 66 to 5 years). For Roseau, two gradients can be identified: a shallow one (from 60 to 7 years) and a steeper one from 7 to 1 year before the eruption. (g,h) evolution of the magmas prior to the large explosive eruption with magma mixing following the period of times shown in (f). T: temperature; black arrow: magma mixing. SZ: single-zoned opx and MZ: multiple-zoned opx. This Figure has been drawn using Microsoft Office suite 2019 Version 1808 (https:// www. micro soft. com/ fr-fr/ micro soft-365).

  the 355 µm fraction (same abbreviations as in Figure2). This Figure has been drawn using Microsoft Office suite 2019 Version 1808 (https://www.microsoft.com/fr-fr/microsoft-365).zoned opx from PPR3, with EPMA profiles indicated with a green line; cores, inner and outer rims are represented (i: inner rim and o: outer rim). a. single-zoned opx; b-c: multiple-zoned opx; d-f: EPMA profiles acquired in a-c in terms of En content (%) with the different plateaus in green that can be linked to different storage conditions corresponding to different ME. d. the plateaus (core then outer rim -first plateau) correspond to the two ME identified for PPR3: ME1PPR3: En52-57 and ME2 PPR3: En48-52; e. the plateaus (core then inner rim -first plateau and outer rimsecond plateau) correspond to ME identified for PPR3 and the same as in d: ME1PPR3 = En52-57 and then ME2PPR3 = En48-52 then back to ME1PPR3 = En52-57;

Figure

  Figure has been drawn using Microsoft Office suite 2019 Version 1808 (https://www.microsoft.com/fr-fr/microsoft-365).

3

 3 magma volumes in the two ME for each eruption, estimated by the proportions of unzoned and zoned opx (on the one hand, from the proportions of Figure 2 and on the other hand, from the histograms in Figure

2

  Institut des Sciences de la Terre de Paris (ISTeP), UMR 7193, CNRS-Sorbonne Université, Paris, France First draft. Aimed journal for the submission, after revisions of the draft: Journal of Volcanology and Geothermal Research.

  . These volcanoes are mainly of three types: either mainly basaltic (Klyuchevskoy, Ushkovsky, Tolbachik), andesitic dacitic (Zimina) or purely andesitic (Bezymianny, the volcano of interest (Figure V.1a-b), and Udina) (Dobretsov et al., 2012).

Figure V. 1 :

 1 Figure V.1: Bezymianny volcano in Kamchatka (summit location marked with a triangle). a) Kamchatka map with the three volcanic zones: Sredinny Range (SR), Central Kamchatka Depression (CKD) and Eastern Volcanic Front (EVF) (Churikova et al., 2001; Melekestsev et al., 1995). Image credit: Nasa Earth Observatory image from January 18, 2017, captured by the Moderate Resolution Imaging Spectroradiometer (MODIS) on NASA's Aqua satellite; courtesy of Jeff Schmaltz. Inset shows the location of Kamchatka on the globe (credit: Google Earth V

  3): logD = -5.54 + 2.6X 𝐹𝑒 -12530 𝑇 (V. 3)

Figure V. 2 :

 2 Figure V.2: Total Alkali Silica (TAS) diagram of Bezymianny. Compositions are shown for whole rocks, melt inclusions and residual glasses for the 1956 blast and Plinian phase (Supplementary Data V.1). Whole rocks compositions of the Plinian phase are from d'Augustin (2021) (Supplementary Data V.1) and literature data for Bezymianny are shown from 1956 to 2010 (Turner et al., 2013). The composition of the preclimactic ashes (Vulcanian-type ash eruptions) are specified (collected 12 km away from the volcano on 7 November 1955, with opx, plag and small inclusions of mgt), as well as blast deposits from the 1956 climactic phase (different lavas than the preclimactic phase, more acid amph andesites) and a lava dome fragment from the new

  3 and 13-16 wt% (Supplementary Figure V.7). Plinian phase MI in plag, amph, opx and mgt range from 72.7 to 82 wt% SiO2, 2.8-7.3 wt% Na2O+K2O, with richer alkaline contents for MI hosted in amph (Figure V.2; Supplementary Data V.1). The residual glasses are rhyolitic with 75.8-81.8 wt% SiO2 and 4-6.3 wt% Na2O+K2O (Figure V.2; Supplementary Data V.1). The compositions of the MI and residual glasses are in the same domain in the K2O, Al2O3, CaO and FeO versus SiO2 diagrams (Supplementary Figure V.7). MI hosted in zoned opx have CaO and Al2O3 contents between 1.6-2.9 and 11-14 wt% and the En content of the hosted opx are between En60-67 (Supplementary Figure V.7).

2 :

 2 Temperatures estimated using magnetite/ilmenite pairs (Ghiorso and Evans, 2008; Sauerzapf et al., 2008; Yavuz, 2021). The temperatures estimated in opx hosted MI from the Plinian phase are also specified (Putirka, 2008).

Figure V. 3 ;

 3 Supplementary Figure V.8-9). The multiple-zoned opx had 2 to 5 boundaries or zonations and sometimes more than 5, with oscillating bands (Figure V.3-4; Supplementary Figure V.8-9). Some opx also show Al sector zoning (Supplementary Figure V.4).

Figure V. 3 :

 3 Figure V.3: Schematics of zoning types identified in the opx of Bezymianny 1956. a) Unzoned opx; b-c) Two types of single-zoned opx: normal-zoned (b), with a Mg-rich core (darker zone on a grayscale image) and Fe-rich rims (clearer in grayscale), or reverse-zoned (c), with inversely a Fe-rich core and Mg-rich rims; (d-i) Multiple-zoned opx, either normal + reverse (d) or reverse + normal (e) or with three rims (normal + reverse + normal (f), or reverse + normal + reverse rims (g)) to four (h) or more than five bands (i). BSE images corresponding to these schemes are shown in Supplementary Figure V.8. White areas are mgt crystals or inclusions. j) BSE image of a multiple-zoned and unzoned opx with the corresponding Red-Green-Blue (RGB) chemical map displaying Mg and Fe elements on (k). Mg is in blue and Fe in red.

Figure V. 4 :

 4 Figure V.4: Proportions of the zonations identified in the opx of: (a) the blast in all size fractions (500-710, 315-500, 250-355 and 125-250 µm) and (b) the Plinian phase, in the size fractions between 315-500, 250-355 and 125-250 µm. The major zonations are highlighted in grey. The detailed proportions of zonations in each fraction are shown in Supplementary Figure V.9. NZ: normal-zoned opx, RZ: reverse-zoned and MZ: multiple-zoned.

Figure V. 5 :

 5 Figure V.5: En contents of the zoned and unzoned opx of the blast and Plinian phase. Probability density functions (PDF) of the En content of unzoned (a-b) and zoned opx (rims and cores) (c-d). Gaussian Kernel Density Estimation (KDE) to estimate the probability density functions of the En content of the opx adapting the code of Petrelli (2021). e-f) Reverse and normal last rims (N) are shown, as well as the boundaries before (N-1) and their zoning types compared to the N-2 boundary (either normal or reverse). Single-zoned rims also shown for opx of the Plinian phase (dark dashed histograms). An identification of two magmatic environments can be shown, with rims evolving either towards En-richer or En-poorer contents (e-f), separating the En content of the opx between En58-65 and En65-68 (ME2 and ME1, respectively). n is the number of opx analysed.

Figure V. 7 :

 7 Figure V.7: Compositions of the mgt (a-b) and curves of the timescales modelled in the blast and Plinian phase (c-d). a-b) TiO2 vs FeO compositions of the zoned mgt for the blast and Plinian phase (for single-zoned ones (cores and rims) or multiple-zoned ones (cores, inner rims and outer rims; one mgt for the blast and two for the Plinian phase)); c-d) Timescales modelled in the reverse-zoned mgt. The main manifestations of the preclimactic volcanic activity prior to the flank collapse and laterally directed blast shown on Figure V.6 are also shown here (from the end of September 1955 to 30 March 1956) (Belousov and Belousova, 1998; Belousov and Bogoyavlenskaya, 1988; Gorshkov and Bogoyavlenskaya, 1965). Uncertainties on individual timescales calculations are shown in Supplementary Figure V.13 and Supplementary Data V.5. n in (a-b) is the number of mgt investigated and n in (c-d) is the number of timescales modelled.

  V.4a). These zonations are associated with En changes characterized by magmatic environments changes (ME1 to ME2 and ME2 to ME1; with ME1 = En65-68 and ME2 = En58-64) (Figure V.5; Supplementary Figure V.11).The timescales modelled on all the rims (reverse or normal zonings) are mainly below one year. Up to 3 years before the blast, normal zoning was mainly produced, as well as some reverse zoning in inner rims of MZ opx (Figure V.6a). However, only one timescale recorded this long timescale of 3 years, so the representativity of this timescale may be questioned. The inner zonations may have been produced by magma convection in the reservoir at ~9-15 km depth, either forming normal zoned opx by mixing with colder areas or reverse zoning with hotter areas(Couch et al., 2001). The N-1 normal-zoned rims, mainly 6 months before the climactic phase (Figure V.6ab), could correspond to magma decompression during the slow magma ascent to form the cryptodome.

  Lange, 2017) (Figure V.6a-b; see Supplementary Note on Fe 2+ -Mg exchange coefficient KD change and H2O content change). The reversely-zoned mgt also indicate a recent heating or changes in fO2, shortly before the lateral blast (mainly below one month; Figure V.7). Plinian phase MZ opx are also dominant in the Plinian phase, with the same compositions as opx from the blast and the same ME evidenced (Figure V.4b; Figure V.5). Timescales are all shorter than a year (Figure V.6c).

  V.6c). These normal zonations could then correspond to a beginning of decompression of the magma of the Plinian phase (Figure V.6c). Then, mainly 10 days before the eruption, reverse zoning is formed in the opx and in the mgt, showing some heating of the magma, possibly due to a magma injection at depth and/or degassing processes (Figure V.6c; Figure V.7).

  .15)), resorbed plag (suggesting a magma injection[START_REF] Ustunisik | New insights into the processes controlling compositional zoning in plagioclase[END_REF]) and reverse zonations in opx and mgt(Figure V.9d). Degassing of the magma or magma injection from below could explain these zonations. Furthermore, open-system outgassing took place within the cryptodome, showed by a density study of the 1956 cryptodome deposits[START_REF] Neill | Influence of pre-eruptive degassing and crystallization on the juvenile products of laterally directed volcanic explosions[END_REF]. As the cryptodome had a long time-building period, it led to crystallization and degassing, as evidenced in the residual glasses of the blast with high microcrystallization (Figure V.2). Reaction rims in amph in cryptodome samples are thought to have been developed under isobaric heating following a heating process, less than 37 days before the blast[START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF] (Figure V.8). These amph can have recorded the heating event in the reservoir and correspond to the last magmas injected in the cryptodome, forming the less degazed core of the cryptodome[START_REF] Neill | Influence of pre-eruptive degassing and crystallization on the juvenile products of laterally directed volcanic explosions[END_REF][START_REF] Shcherbakov | Phase equilibria constraints on preeruptive magma storage conditions for the 1956 eruption of Bezymianny Volcano, Kamchatka, Russia[END_REF] (Figure V.8). 5) For the magma giving the Plinian phase, less than 6 months before the climactic phase, decompression of the reservoir in relation with the beginning of ascent of a part of the magma in the cryptodome led to normal zoned rims in opx (Figure V.9b-c). Some reverse zoning is also produced due to heating from below. Then, approximately one month before the eruption, significant normal zonations show the decompression of the reservoir (Figure V.9d). Finally, less than 10 days before the climactic phase, the crystals record a heating process or degassing (Figure V.9e).

. 4 ;

 4 Supplementary Figure V.7).From an end-member analysis,[START_REF] Turner | The magma plumbing system of Bezymianny Volcano: Insights from a 54year time series of trace element whole-rock geochemistry and amphibole compositions[END_REF] also inferred that a magma injection occurred from depth (> 30 km), showing a magma recharge in the deeper reservoir(Turner et al., 2013) (Figure V.8). They also evidenced that, as only this shallow reservoir is connected to the vent, the eruption of 1956 only emits magma from the upper part of the shallower reservoir. 6) Then, collapse of part of the flank of the volcano led to the decompression of the cryptodome, generating the lateral blast and leading to the decompression of the shallow reservoir below, inducing the fast ascent of the magma at the origin of the Plinian phase (Figure V.9f). A Plinian column was generated, giving ash and pumice fallout and pyroclastic density currents by collapse of the eruptive column.

Figure V. 8 :

 8 Figure V.8: Summary of the crystal assemblages of the blast and Plinian phase, crystal compositions, storage conditions and dynamics of magma storage inferred from crystals zoning (opx or mgt; this study), resorption zones in plag (Shcherbakov et al., 2013) or reaction rims in amph (Plechov

Figure V. 9 :

 9 Figure V.9: Magma dynamics inferred for the 1956 climactic phase of Bezymianny, from the combined point of view of opx and mgt of this study and plag and amph from cryptodome clasts and pyroclastic deposits from the Plinian phase (see Figure V.8). a) Pre-1955 (up to 3 years for opx), magma

  L. O., G. B., H. B.B., A. B., M. B. and T. d'A. participated in the field mission in 2019 to have access to the samples in Kamchatka (Russia). L.O. performed data acquisition, prepared the figures and wrote the initial version of the manuscript during herPhD. L. O., G. B. and H. B.B. participated in the interpretation of the data and the writing of the manuscript text.

Figure

  timescales results. Median of the timescales in opx, expected timescale and its standard error associated for each dataset are shown, according to the changes of some parameters (T, fO2 or not, D), for the blast and Plinian phase, as shown in Supplementary the temperature found using Sauerzapf et al.(

D

  'après la modélisation de l'interdiffusion intracristalline des éléments Fe-Mg dans 88 cristaux d'orthopyroxènes zonés, les magmas ont subi des phénomènes de mélanges des magmas ~1,5 ans avant l'éruption, qui sont corrélés avec le début d'une crise sismique (Chapter VI.1). Ceci pourrait fournir, sur les volcans d'arc, une information précieuse pour l'interprétation des éruptions passées monitorées et futures en cas de réactivation. Combinant l'étude les cristaux de plagioclase et de magnétites zonés contenus dans les mêmes échantillons que les orthopyroxènes, les plagioclases enregistrent un événement d'injection de magma dont la temporalité n'est pas contrainte et qui pourrait avoir eu lieu en 1963 lors d'une crise sismique au Kizimen. Cet événement est suivi par un début de mélange des magmas entre 1 et 5 ans avant l'éruption, enregistrés par quelques plagioclases et orthopyroxènes, et pouvant correspondre à la faible activité sismique entre 2005 et 2008. Ensuite, ~1,5 ans avant l'éruption jusqu'à quelques jours avant l'éruption, un mélange plus vigoureux des magmas andésitiques et dacitiques a lieu, enregistré par les orthopyroxènes, plagioclases et magnétites (Chapter VI.2). Ces événements sont corrélés à la crise sismique débutant en avril 2009. L'étude combinée de ces trois types de cristaux permet d'avoir un modèle plus complet des dynamiques pré-éruptives des magmas, en termes de processus et d'échelles de temps.

  In the form of the accepted manuscript in October 2022 in Communications, Earth & Environment. It is presented here with numbered lines and with the original titles without numbers, as submitted to the journal. Published form accessible with this link: https://doi.org/10.1038/s43247-022-00622-3 Pre-print link on Research Square: https://doi.org/10.21203/rs.3.rs-1140681/v1

1 ;Figure 3 ). 2

 132 Figure3). Zoned crystals were classified according to the type of zonation observed: either a single one (normal or reverse zonation) or a multiple one (Figure1g-j; Supplementary Figure2-3). In the dacite, the majority of the zoned opx are normal-zoned while, in the andesite, reverse-zoned opx are more abundant (Supplementary Figure2-3). Few opx show reactional borders or remelting textures (Supplementary Figure2-3). Out of the 1841 opx crystals mounted, 112 zoned opx compositions and 108 unzoned opx compositions were investigated in the dacite and andesite (Figure2; Supplementary Data 2; Supplementary Figure4-5). Opx compositions are spread over an En44-73 compositional range (Enstatite content (𝐸𝑛 = 𝑀𝑔 𝑀𝑔+𝐹𝑒

Figure 5 .

 5 Compilationof all the data on these studies correlating these two types of information was done to reflect on other studies linking petrological data and monitoring signals (Supplementary Data 5). Kizimen timescales are in good agreement with other pre-eruptive timescales modelled in basaltic andesitic or dacitic volcanoes, such as Ruapehu (1969 eruption), Bezymianny (2012 eruption), which is also in Kamchatka, Llaima (2008 eruption) and the 2014 eruption of Piton de la Fournaise (basaltic) (Figure5; Supplementary Data 5). The number of timescales modelled here is high, compared to the other studies and the relationship with the monitoring signals is well correlated (Figure5; Supplementary Data 5).

  Figure 1. From macroscale to microscale: Kizimen location in Kamchatka, 2010-2013 eruption samples and opx zonations. a) Kamchatka map with the three volcanic zones: Sredinny Range (SR), Central Kamchatka Depression (CKD) and Eastern Volcanic Front (EVF) 33,34 and the boundary of the active subduction zone of the Kurile-Kamchatka Trench (subduction rate of 8

(

  alternation between dacitic and andesitic bands 41 ). On the field, estimates of the percentages of these different samples were ~70% of banded lavas, ~10% of dacites and ~20% of andesites (even if it is unclear how the initial products looked like or what fraction remained in the chamber). gh) Backscattered electron images (BSE) of major zonations identified in the dacite and andesite: normal and reverse zonations, respectively. i) Red-Green-Blue (RGB) chemical map of a normalzoned opx, with a Mg-rich core and Fe-rich rims ("Outer" on (i-j)). Mg is in blue and Fe in red. j) Reverse-zoned opx, with a Fe-rich core and Mg-rich rims, with the same colors as in Figure 1i.

Figure 2 .

 2 Figure 2. Enstatite (En) content histograms of unzoned (a) and single-zoned opx of the 2010-2013 eruption (b) (unzoned opx compositions shown in the dashed frame). The arrows underline the major path from the core to the rim recorded by the single-zoned opx. The number of opx analyzed is specified (n). The cores of the zoned opx are in the same range as the unzoned opx.

Figure 3 .

 3 Figure 3. Linking monitoring data with timescales of magmatic processes. a) Time vs depth plot of seismicity at Kizimen from 2003 to the onset of the eruption (11/11/2010) 22,41 . The onset of the surface deformation detected by InSAR is specified 24,41 . b) Cumulative number of the timescales estimated on the major compositional changes evidenced for the dacite and andesite (normal zoning for the dacite (En62-64 to En44-54) and reverse zoning for the andesite (En62-64 to En64-73); Supplementary Data 4) and cumulative seismic moment. c) and d) Focus on the 1.5 years preeruption. More crystal clocks are started closer in time to the eruption (steeper slope one year before the eruption). e) Uncertainties associated to the individual timescales modelled shown on c) (Supplementary Figure6, Supplementary Data 4). These uncertainties are due to the diffusivity calculations (D0, the interdiffusion coefficient, the activation energy Ea), temperature and resolution of grayscale values of the BSE images6,112 . They are calculated using a Monte Carlo simulation. These uncertainties are asymmetric, with a larger positive error bar and a smaller negative error bar, due to the logarithmic effect of the uncertainties on mainly temperature and diffusivity calculations26 .

Figure 4 .

 4 Figure 4. Pre-eruptive dynamics of the magmas in the timeframe given by the timescales modelled in zoned opx: crystal clocks mainly starting 1.5 years before the eruption. a) Magma plumbing system below Kizimen before the 2010 eruption. The architecture of the plumbing system of a previous study was used 41 : a dacitic reservoir which has been intruded by a basaltic injection, possibly in 1963 22,25,41 , with the lower part of the dacitic reservoir becoming andesitic by hybridization with the mafic magma. During this hybridization, crystals of the dacite get mixed with the andesite (cores of all opx are the same as the dacite) (Figure 2; Supplementary Figure 4-5). Then, the most likely hypothesis to explain opx zonations in the dacite and andesite is presented (see Discussion): b) Magma mixing occurred 1.5 years before the eruption, at the contact between the dacite and the andesite, exchanging opx from the dacite with the andesite (explaining the formations of reverse-zoned opx in the andesite by increase of the temperature (T)). Opx crystals can also return in the dacitic magma, forming normal zonations and get reincorporated in the andesite before the eruption (forming multiple-zoned crystals). As a result of mixing of magma with colder areas in the dacite, the opx in the dacite mainly record normal zonations. Not all the reservoir was affected by the mixing process as there is a majority of unzoned opx. Seismicity is recorded during this period. The rate of crystal diffusion starting rates (progressive increase of timescales from April 2009 to the eruption) and progressive increase of seismic moment and number of earthquakes are shown with the shading in the arrows. The deeper part of the transcrustal system may be composed of mafic lenses. The dacitic and andesitic reservoir have the same mineral assemblage: plagioclase, opx, amphiboles, Fe-Ti oxides and some resorbed olivines and quartz (T = 832 ± 37°C and 850 ± 57°C, respectively).

Figure 5 .

 5 Figure 5. Database showing the comparison between diffusion timescales for studies correlating diffusion timescales and monitoring signals before or during an eruption. Most of them involve olivines or pyroxenes (orthopyroxenes or clinopyroxenes). The range of timescales that can be modelled with diffusion methods go from hours to centuries 7 but here, they are mainly between months to a decade before the eruption. Studies with well-correlated diffusion timescales and preeruptive seismic data are specified. Pre-eruptive timescales before the 2010 eruption of Kizimen studied here are shown (for andesite (and) and dacite (dac); see arrows). The studies used for comparison are: Ruapehu (1969, 1971, 1977, 1995, 1996) 20 ; El Hierro (2011-2012) 2,73 ; Eyjafjallajökull (2010) 14 ; Bezymianny (2006-2012) 21 ; Etna (1991-1993 ; 2006)13,15 ; Piton de La Fournaise (2007 ; 2014-2015)71,72 ; Kelud (1990)18 , Llaima (2008)76 ; Shishaldin (1999)16 ; Mt St Helens (1980-1986) 17 ; Kīlauea (1959) 70 ; Merapi (2006 ; 2010) 74 ; Irazú (1963-1965) 75 . Symbols correspond to different minerals and the colors to different magma compositions. The symbol position represents the pre-eruptive timescales given by the minerals (median of the dataset, when

  Kizimen is the southernmost active volcano of the CKD(Auer et al., 2018) (55.131N,160.320E, 2375 m above sea level; south of the Klyuchevskoy group of volcanoes; FigureVI.2.1a). It is located on the south-east flank of the Shchapinsky graben, within a fault system including the junction of the graben with the Tumrok ridge horst(Auer et al., 2018;[START_REF] Shantser | Kizimen volcano[END_REF]. This fault system intersects pre-Kizimen deposits and the volcano itself, such as the SW-NE normal fault located on the northwestern slope of the volcano (Figure VI.2.1b)(Auer et al., 2018).

Figure VI. 2 . 1 :

 21 Figure VI.2.1: Kizimen volcano (Kamchatka). a) Topographic map of Kamchatka (Shuttle Radar Topography Mission (SRTM), NASA/JPL/NIMA picture (taken in February 2000) with the three

  Figure VI.2.1c-d), andesites (Figure VI.2.1e-f) and banded samples (Figure VI.2.1g) (light bands of dacitic composition and dark andesitic bands (Auer et al., 2018). Banded lavas often appear with bands from a few millimeters to 20-25 cm wide (Figure VI.2.1g).

  e. the distribution of Mg values in the crystal before the beginning of diffusion that changed the Mg profile. The estimation of the initial Mg profile is based on the form of the Mg profile obtained with EPMA. The initial Mg profile is a step profile with high and low plateaus defined by the maximum and minimum Mg values between the rims and the intermediate zones. By multiplying the acquisitions of profiles in the same crystal, it increases the number of analysed points and improves the shape of this initial Mg profile. Moreover, we can estimate the initial Mg profile using the chemical compositions of An and the partition coefficients K (𝐾 𝑝𝑙𝑔/𝑙𝑖𝑞 = 𝐶 𝑖 𝑝𝑙𝑔 𝐶 𝑖 𝑙𝑖𝑞

  with a = -26.1± 1.1 kJ and b = -25.7± 0.7 kJ. -The diffusive model can be presented in one or two steps if we consider the crystal in isothermal conditions (a constant temperature T) or in the case where the crystal passes through two successive temperatures Tcore and Trim (a temperature for the formation of the core and a second temperature corresponding to the time of the rim formation) (Figure VI.2.2a-b).The following method is the one relative to a one-dimensional diffusion model, under isothermal conditions. The plagioclase diffusion model code was adapted from the code available in the PhD thesis of Fabbro (2014) and can be available upon request. zonation in andesite; g) Normal-zoned rim with a "patched" core in the andesite; h) Oscillatory zonations in dacite. The four types of plag are specified; type 4 is only found in dacitic samples. Mgt: i) Zoned mgt; j) Exsolved mgt. The white and black transects in crystals correspond to the diffusion profiles carried out with EPMA. Plag: All observed plag are zoned. The plag observed in the samples were divided into four types according to the major zonation identified on the rims of the crystals or remarkable characteristics (resorption surfaces, well individualized cores, patched zones; Figure VI.2.3): • Type 1: defined by a reverse-zoned rim (called RZ hereafter; An increase from core to rim). The rim (~10-20 µm thick) is easily identifiable because it is very clear in the SEM images. It also appears very rich in Ca on the chemical maps (Figure VI.2.3d). It is preceded by a sawtoothlike resorption surface (Figure VI.2.3e). • Type 2: This type of plag shows a normal zonation at the outermost rim (called NZ hereafter; An decrease from core to rim). This type is recognizable by a variation of shades of gray going from clear (calcic) in intermediary zone to darker zones (less calcic) toward the edge (Figure VI.2.3f). This change can be either gradual or abrupt. • Type 3: The core is, in rare cases, partially or totally resorbed. This important resorption is called the "patched zone" (Figure VI.2.3g). • Type 4: It shows oscillatory zonations corresponding to many An variations (of 5%) (alternation of multiple and fine zonations close together (from 2 to 5 µm)) (Figure VI.2.

Figure VI. 2 . 4 :

 24 Figure VI.2.4: Histograms of compositions measured in opx and plag of the dacite (shades of blue) or andesite (shades of red). Opx: a) Cores and (e) rims of opx compositions (En (%)) of dacites and andesites (all data from Ostorero et al. (accepted)); Plag: b) Cores compositions (An (%)) of all plag; c-d) Intermediate zones of the plag of the andesite (25 % recording a low An content and 75 % recording higher An content); f-g) Outermost rims of the plag, with (f) 100 % of the rims of the dacitic plag. Arrows 1 and 2 mean different magmatic processes defined by variations of An/En contents in the discussion. n is the number of crystals analysed in the dacite or andesite.

Figure VI. 2 . 5 :

 25 Figure VI.2.5: Mgt composition and plag compositions in minor elements of the dacite (shades of blue) or andesite (shades of red). a) TiO2 vs FeO compositions of the zoned mgt (cores and rims) for the andesite and dacite. b) Fe vs An (%); Fe trace element signals in plag. The boxes highlight the major areas of plag compositions depending on the dacitic or andesitic nature of the sample.

Figure VI. 2 . 6 :

 26 Figure VI.2.6: Histograms of the timescales after modelling intracrystalline diffusion. a) Interdiffusion modelling of Fe-Mg in opx (dacite in blue and andesite in red) (all data from Ostorero et al. (accepted)); (b) Mg in plag and Ti in mgt with different colours corresponding to the plag types and zonations types in mgt (2 timescales on inner to outer rims in multiple-zoned mgt (MZ)). n is the number of crystals analysed in the dacite or andesite.

Figure VI. 2 . 7 :

 27 Figure VI.2.7: Uncertainties associated to the timescales in plag and mgt. The uncertainties are due to the uncertainty on the estimated temperatures: 847 ± 57 °C for the andesite in plag (a) and mgt and 832 ± 37°C for modelling on mgt of the dacite (b). For uncertainties associated to opx timescales, see Ostorero et al. (accepted).

2 . 2 :

 22 Timescales modelled in plag of the andesite. The name of the crystals, zonation types (NZ: normal-zoned, RZ: reverse-zoned, RC: resorbed core and errors associated to the plag are specified (corresponding to the modelling with the uncertainties on the temperature considered). An changes where the timescales have been modelled are also specified as An ranges are the An variations between the two plateaus are important.

Figure

Figure VI. 2 . 8 :

 28 Figure VI.2.8: Crystal System Analysis approach of the zonations recorded by the plag and opx for the dacite (a-b) and for the andesite (c-d). Each box represents a magmatic environment (ME): ME1 = An42-54, En62-64; ME2= An72-84 for plag, ME3 = An60-70, En64-73 and ME4 = An30-50 and En42-62. For opx, the sizes of the boxes are approximately proportional to the proportions of unzoned and cores of zoned crystals in each ME. The relative sizes of ME1 are similar for the dacite and andesite, as the crystallinity is the same (~99% in volume, as all unzoned opx and practically all the cores of the zoned opx are in this ME1). For plag (a and c), the change between ME recorded from the core to the rims of the crystals are represented by a black arrow, while a dashed arrow represents a resorption. For opx (b and d), the change of ME recorded from the core to the rim of opx is represented with a black arrow for single-zoned crystals and a dashed arrow for multiple-zoned crystals. The percentages of crystals that record the particular ME changes are specified next to the arrows. For opx, 80 % of the timescales estimated are characterized by a change from ME1 En62-64 → ME3 En64-73 for the andesite and ME1 En62-64 → ME4 En42-62 for the dacite. For plag in the andesite (c), 40 % of the timescales estimated are characterized by a change from ME1 An42-54 → ME2 An72-84 and 60 % of them correspond to a change from ME1 An42-54 → ME3 An60-70. No timescale is estimated for plag from dacite. See Figure VI.2.6 and Table VI.2.2 for plag timescales.

  Figure VI.2.4f). These compositions describe a new magmatic environment ME3: 25% of plag pass directly from ME1 to ME3 (Figure VI.2.8c). These plag show reverse zonations preceded by a resorption surface with a sawtooth pattern (Figure VI.2.3d-e; Table VI.2.1, Figure VI.2.8c). The majority of opx also record significant changes from ME1 En62-64 to ME3 En64-73 (65 %; Figure VI.2.8d).

Figure VI. 2 .

 2 8b; d). This could occur in the dacitic part of the reservoir, following some mixing with colder areas (Figure VI.2.8b; d).

  Moving from the CSA to the timescales, the diffusion timescales modelled in plag and opx correspond to the entry of the crystals into ME3 (Figure VI.2.6b; Figure VI.2.8c). ME3 is the result of interactions and mixing between the remnant dacitic and newly formed andesitic magma (Figure VI.2.8b-d).

  years prior eruption up to the eruption (Figure VI.2.8b; d) and also by plag (Figure VI.2.8a; c). Indeed, plag are integrated in ME3 where they crystallize the An60-70 rims at two different times: -Between 2 and 5 years for type 2 and 3 of plag, after an injection of mafic magma which is not constrained in time (Figure VI.2.8c). -Within 6 months before eruption for type 1 of plag that were originally derived from dacitic magma (Figure VI.2.8c). These timescales are superimposed on those of the opx (Figure VI.2.8b; d).

Figure VI. 2 . 9 :

 29 Figure VI.2.9: Cumulative number of the timescales estimated on the major compositional changes evidenced in plag, opx and mgt. The timescales are modelled in plag in the andesite (type 1; reverse zoning (An42-52 to An60-70) and normal zoning (types 2 and 3; An78-84 to An62-75)) (Figure VI.2.6, Table VI.2.2), for the opx in the dacite and andesite (normal zoning for the dacite (En62-64 to En44-54) and reverse zoning for the andesite (En62-64 to En64-73) (from Ostorero et al. (accepted)) and mgt (reverse-zoned and normal-zoned rims in andesite and dacite). b) Focus on the 1.5 years pre-eruption. c) Focus on the 3 months pre-eruption. The seismic crisis period is specified in (ac) (Senyukov et al., 2011; Shakirova and Firstov, 2022).

Figure VI. 2 . 10 :

 210 Figure VI.2.10: Model of the reservoir dynamics during the recent history of Kizimen prior to the 2010 eruption recorded by plag, opx and mgt, on different timescales.

  Plag and opx indicate that all crystals initially begun to crystallize in the dacitic magma present as a single reservoir (ME1)(Figure VI.2.10a). More than five years before the eruption (probably around 1963, correlated with the seismic crisis(Auer et al., 2018;[START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF]), an aphyric mafic magma (absence of crystals with clearly mafic composition) from a deeper source is injected at the base of the reservoir (Figure VI.2.10b). This recharge event heats part of the system and causes the melting of some of the plag and few opx caught up in this new melt (Figure VI.2.10b). Only a part of the reservoir, probably the lower part of the reservoir, is affected by the injection which leads to the mixing between the mafic magma and the dacite (ME2). This results in the formation of a temperature-zoned reservoir with an andesitic magma at the base, denser (ME2) and a dacitic magma at the top, lighter (ME1) (Figure VI.2.10c). No recharge is observed in the months preceding the eruption. The plag rims that grew after these resorbed zones recorded long timescales up to 5 years before eruption.At the contact between the two magmas, a mixing process occurred, 1-5 years before the eruption (Figure VI.2.10e) but is mainly recorded 1.5 years prior eruption by opx. This leads to the incorporation of opx from the dacite in the andesite reverse-zoned opx) and a cooling of the dacitic magma at the top of the reservoir (Figure VI.2.10e

Figure VI. 3 . 1 :

 31 Figure VI.3.1: Kamchatka peninsula and Kizimen volcano. a) Topographic map of Kamchatka (Shuttle Radar Topography Mission (SRTM), NASA/JPL/NIMA picture (taken in February 2000) with the three volcanic zones (the Eastern Volcanic Front (EVF), the Sredinny Ridge (SR) and the Central Kamchakta Depression (CKD) (Churikova et al., 2001; Kayzar et al., 2014) and location of Kizimen in the CKD, one of the largest and most active cluster of subduction volcanoes (Journeau et al., 2022; Ponomareva et al., 2007) (55.131N,160.320E, 2375 m asl; south of the Klyuchevskoy group of volcanoes, including Bezymianny). This volcanic arc results from the subduction of the northeast corner of the Pacific plate at a rate of ~8 cm/yr (DeMets, 1992) beneath the Okhotsk plate (which is connected to the North American plate) (Gorbatov et al., 1997; Park et al., 2002). The 2010-2013 eruption of Kizimen emitted 0.5 km 3 Dense Rock Equivalent (DRE) of magma (Auer et al., 2018; Melnikov et al., 2011), with explosive episodes that generated block and ash pyroclastic density currents followed by the extrusion of a lava dome that gradually formed a very thick lava flow (b) (Auer et al., 2018; Churikova et al., 2007;

  Dacitic and banded andesites were collected in the pyroclastic deposits of the 2010-2013 eruption (Figure VI.3.1c; Figure VI.3.2a-b) during a three-week field campaign in Kamchatka in 2019. The coordinates of the samples can be found in Ostorero et al. (accepted). Thins sections in the natural samples were done in the different fragments at the workshop and analytical platforms of Sorbonne Université, Institut des Sciences de la Terre de Paris (ISTeP). The thin sections are used to observe the textures and crystals under the optical microscope and perform textural and chemical analyses described below.

3 . 1 :

 31 Major elements composition of the basaltic enclave of Kizimen (Churikova et al., 2013b) that was reproduced through the synthetic basalt and the starting material compositions of the synthetic basalt and dacite, with the positive and negative errors.

Figure VI. 3 . 2 :

 32 Figure VI.3.2: Andesite-dacite mingling textures of Kizimen 2010-2013 eruption samples shown at different magnifications, on a-b) macroscopic scale and c-d) backscatter electron (BSE) images.

Figure VI. 3 . 3 :

 33 Figure VI.3.3: Schematic representation of the COMMA (not to scale, modified from Morgavi et al., 2015; Rossi et al., 2017). a) Experimental setup showing the high-temperature furnace, the upper and lower motors for the rotation of the spindle and the crucible with the two end-member melts. b) Focus on the experimental crucible with the dacitic composition where a glassy cylinder of basaltic composition is inserted. The off-center position of the spindle is also shown. c) Picture of the crucible, with the basaltic cylinder and hole for the spindle. d) Top-view of the general geometry of the COMMA to perform the chaotic mixing of high-viscosity silicate melts at high temperatures (1200 °C), inspired by the Journal Bearing System geometry (JBS; Swanson and Ottino (1990)). The main experimental parameters are Rout (18.5 mm) radius of the outer cylinder with the dacitic melt, Rin (6.2 mm) radius of the inner cylinder sheated with the spindle, Rb radius of the basaltic melt (5.5 mm), Ωout (0.11 rotation per minute; rpm) angular velocity of the outer cylinder, Ωin (0.66 rpm) angular velocity of the inner cylinder (Morgavi et al., 2015; Rossi et al., 2017).

  (Figure VI.3.4). The same crystal assemblage is present in the different bands of the banded lavas (plagioclase, amphibole, orthopyroxene, Fe-Ti oxides, as well as trace amounts of olivine, quartz and apatite in the different bands)(Auer et al., 2018), while the matrix glass is different between the different bands (FigureVI.3.2c-d).Banding is mainly caused by the different textures of the matrix of the andesitic and dacitic glassy domains, with the andesitic domains that are richer in microphenocrysts than the dacitic glass which is almost free of microphenocrysts and more vesiculated (Figure VI.3.2c-d). These mixing features provide us with unique information about the timescales of eruption triggering as the crystallization of the matrix corresponds to the last phases before the eruption (magma mingling and ascent).

Figure VI. 3 . 4 :

 34 Figure VI.3.4: Total Alkali Silica (TAS) diagram of all compositions: residual glasses in the crystal enclaves (cores, middle, margins and exterior areas), as well as experimental felsic enclaves present in the crystallized enclaves, transects with or without crystals and end-members compositions. The whole-rocks of the dacites and andesites of the 2010-2013 eruption of Kizimen and residual glasses in the natural samples are also shown for comparison in the grey domain (Ostorero et al., accepted), noted as an asterisk (*) in the legend). The basaltic enclave which we wanted to reproduce with our synthetic basalt is also shown (Churikova et al., 2013b) (") and data from Auer et al. (2018) is also shown (').

Figure VI. 3 . 5 :

 35 Figure VI.3.5: BSE images of plagioclase and oxides microlites in the experimental sample in the basaltic area (a) and natural samples (b-c, in the andesitic and dacitic domains, respectively) and microlite aspect ratios (d). On (d), the ratio of the short axis (S) over the medium axis (M) is shown depending on the ratio of the short axis (S) over the long axis (L). These parameters correspond to the best-matching crystal habit (x y and z, found with CSDSlice (Morgan and Jerram, 2006)). The experimental microlites and dacitic ones are tabular, whereas the andesitic microlites are acicular.

3 . 3 :

 33 CSD results. Estimates of the growth rate G and nucleation rate J of the crystals (mm/h) in the experimental sample and estimates of the residence timescales τ of the microlites in the banded lavas of 2010 of Kizimen, in the andesitic and dacitic bands, using different experimental growth rates and the two different slopes of the CDS curves for the dacitic and andesitic bands in the natural banded andesites (for the andesitic band: using the minimum and maximum experimental growth rates for a basaltic andesite (1

Figure VI. 3 . 6 :

 36 Figure VI.3.6: Crystal Size Distributions (CSD) of the experimental (basaltic area) and natural samples of Kizimen (in the andesitic and dacitic domains), derived by CSDCorrections 1.5 software (Higgins, 2002, 2000). a-b) and estimates of residence timescales of crystals from natural samples, using the two parts of the slopes of the curves of the CSD and different experimental growth rates (Hammer and Rutherford, 2002; Mollard et al., 2012; Shea and Hammer, 2013) (cd). (b) and (d) are zoom on the dashed framed rectangles. The linear best-fit slopes are shown on (a-b) and are reported, with the intercepts of the three CSD in SupplementaryTable VI.3.1. Timescales are reported in Table VI.3.3.

Figure VI. 3 . 7 :

 37 Figure VI.3.7: BSE images of the experimental samples. a) Overview of a representative slice of the experiment, which shows a dark grey background (the felsic domain), 2-3 big bubbles (~1mm) and one wide light grey area (the hybrid domain; where the image of Figure VI.3.7b was taken). The width of the light areas varies between the different crystallized areas but are between 300-2000 µm and can reach 1 cm. The big hole corresponds to the position where the turning spindle was during the experiment. Few other light grey areas are dispersed around the spindle but are smaller in size. The main light grey areas are located in the upper left part and below the spindle. Filaments are connecting these areas together; b) Hybrid crystallized enclave with plagioclase crystals with different sizes and small oxides. An interaction zone is present between the felsic and mafic areas (white dashed line); c) Small crystallized areas are dispersed in the experimental sample (called "crystal enclaves"). Some crystals originated from the main crystallized areas are located close to the small crystallized zones dispersed in the sample or near some filaments but they are isolated in the felsic areas; d) Hybrid filament with a small crystal enclave with plagioclase crystals. The filaments show some deformation with stretched and folded filaments; e) Felsic enclave are present in the margins of the crystallized areas, also showing some isolated crystals.

  (Figure VI.3.7a). The mafic melt generated a filament-like morphology in the felsic melt around the inner cylinder (Figure VI.3.7a). The crystal enclaves are connected to some mafic filaments around the hole of the spindle (Figure VI.3.7a), with crystals in the hybrid areas. As these crystallized portions are evidenced in the hybrid areas but as the glassy end-members before the experiment were homogenized and contained no crystal (Supplementary Figure VI.3.1), the hybrid part thus crystallized during the chaotic mixing experiment. Both mixing and crystallization occurred during this experiment, as the temperature chosen was below the liquidus of the plagioclase of the basalt. In these crystal enclaves, plagioclase crystals of different sizes and small oxides are found (Figure VI.3.7b-d). The size of the crystals varies depending on their location from the margins of hybrid enclaves (Figure VI.3.7b-c; Supplementary Figure VI.3.4), with the crystals in the cores of these enclaves having a higher size (80-300 µm) than the ones in the middle (125-180 µm) and in the margins (~40-50 µm) (Supplementary Figure VI.3.4). In the felsic areas, the isolated crystals are between 20 to 100 µm (Supplementary Figure VI.3.4). Isolated crystals in the felsic areas likely detached themselves from the crystal enclaves during the rotation of the second cylinder and were carried away in the felsic portion (Figure VI.3.7c and e).

Figure

  of the experimental samples are mainly andesitic, resulting from the mixing of the basalt with the dacite, or dacitic, with rare basaltic and basaltic andesite compositions (Figure VI.3.4). Indeed, the transects across the felsic matrix with hybrid filaments show oscillations in the oxide compositions, with areas with lower content in SiO2 but higher content in CaO, MgO and FeO, with the filaments corresponding to a hybrid melt, which results from the mixing (Figure VI.3.4 and 8; Supplementary Figure VI.3.5; Supplementary Note). Around crystal enclaves, an interaction zone between the end-members developed (Figure VI.3.7b-c), with the margins having glass of intermediate compositions between the felsic and hybrid (Supplementary Figure VI.3.5

  e. diffusion coefficients, different degrees of hybridization are produced, but overall, mafic compositions proportions are lower than felsic ones[START_REF] Zhang | Diffusion in Melts and Magmas[END_REF] (Supplementary FigureVI.3.5).

Figure VI. 3 . 8 :

 38 Figure VI.3.8: Compositional variation along three representative transects in the glasses (a1-d3) for major elements: a) SiO2, b) CaO, c) MgO, d) FeO. The Supplementary Figure VI.3.8 shows the locations on BSE images of the transects a1-d1 in the experimental slice. Compositions of initial basaltic and dacitic melts are marked (blue and yellow, respectively), with error bars including analytical uncertainties. a4-d4) Compositional histograms of the compositions of the transects without crystals and with crystals.

  Conversely for large undercoolings, compositional gradients develop in the melt next to the crystal surface. However, most of the crystals are tabular (Figure VI.3.5), some crystals in the experimental samples are dendritic and compositional gradients in the melt next to the crystal surface are found, such as in the Harker diagrams (Figure VI.3.9), where the cores of the enclaves are more mafic than its margins. But no disequilibrium texture is found in the crystals apart from slight normal zoning (2-5 µm) in plagioclase crystals (Figure VI.3.10). This phenomenon is then partially observed in the interaction zones between the end-members (Figure VI.3.7b-c), whereas in the hybrid parts, crystals growth is favored.

Figure VI. 3 . 9 :

 39 Figure VI.3.9: Harker diagrams for the glasses (in crystal enclaves, felsic enclaves, transects) and end-members compositions. a) CaO vs SiO2; b) MgO vs SiO2; c) K2O vs SiO2; d) Na2O vs SiO2; e)

76 (

 76 Figure VI.3.10d; Supplementary Figure VI.3.6). Deformation brings various crystals originating from the mafic component in the felsic areas and they show a sharp contact between the felsic host and their rims, with an An# of 71-73 (Figure VI.3.10a) (Laumonier et al., 2014) and chemical gradients are present in crystal blobs between the felsic parts and cores/middle/margins of these zones (Figure VI.3.9). Experimental plagioclase crystals have a smaller range in anorthite than plagioclase in the banded lavas or from the andesite of the same eruption (An28-86 and An26-70, respectively) (Auer et al., 2018; Dupont de Dinechin, 2021) (Figure VI.3.10; Supplementary Figure VI.3.6). This compositional difference can be due to the fact that their histories are more complex in the natural system.

Figure VI. 3 . 10 :

 310 Figure VI.3.10: Experimental plagioclase crystals and their compositions. a-c) BSE images of plagioclase crystals with numbers that denote their Anorthite content (An#) and d) An# of the crystals depending on their locations (crystal enclaves: cores, middle or margins or in filaments) and in their cores/rims or in the edges of the crystals. Isolated crystals are highlighted on a). The An# predicted by MELTS software for the synthetic basalt at 1200°C and NNO + II is also shown (Gualda et al., 2012).

  Figure VI.3.11a).

Figure VI. 3 . 11 :

 311 Figure VI.3.11: Schematic sketch of the processes deciphered here, the kinetics of the first mixing event related to the injection of a basalt in the dacite, leading to the formation of an andesite (a) and the timescales of the third magma mixing event that occurred very shortly prior to the eruption of Kizimen (few minutes to hours; c). The second magma mixing between the andesite and the dacite is not investigated here (b) (Ostorero et al., accepted).

3 . 5 ).

 35 The glasses in the "middle zone" are in between these two ranges (Supplementary FigureVI.3.4 for the location). These relationships are also visible for the CaO, MgO and FeO contents (Figure VI.3.9b1-d1). The felsic enclaves also sampled some more mafic areas (56 wt% in SiO2) (Figure VI.3.9a2). The basaltic remains are nearly absent whereas dacitic areas are dominant, as well as hybrid zones (Figure VI.3.9a3-d3). In the glasses near the crystallized areas, the basaltic remains are nearly absent whereas dacitic areas are dominant, as well as hybrid zones (Figure VI.3.9a3-d3).Comparing the experimental glasses with the natural ones, the residual glasses from the natural banded andesites are rhyolitic whereas the experimental ones are from basaltic andesites to dacites (Figure VI.3.4). The experimental glasses are closer from the whole rock compositions of the two natural end members, andesites and dacites, as the basaltic and dacitic end-member have been homogenized to do the experiment.VI.3.A.2.4. Chemical compositions of crystal phases In the experimental sample, the crystallized enclaves crystallized during the experiment, as the starting materials were initially crystal-free (Supplementary FigureVI.3.1). The Anorthite contents (𝐴𝑛# = 𝐶𝑎 𝐶𝑎+𝑁𝑎+𝐾

  ). Chemical maps show a slight normal zonation (Na higher on the rims, Supplementary Figure VI.3.7). The An# also show normal zonations both in crystallized enclaves and filaments (<5 µm; Figure VI.3.10a-d, Supplementary Figure VI.3.6) (from calcic cores to less calcic rims). The cores of the crystals are mainly between An60-76 whereas the rims are lower in An#, mainly between An50-63. The edges of the crystals have the same An# as the cores of the crystals (An69-76) (Figure VI.3.10d).

VII. 1 . 2 .

 12 Comparison with pre-eruptive timescales from other volcanoes VII.1.2.1. Diffusion timescales A comparison of the timescales found for our seven eruptions with other eruptions from more than 60 papers from the literature and 86 eruptions where diffusion modelling was carried out is shown in Figure VII.1, either in mafic or evolved systems, for monitored or unmonitored eruptions.

  (Figure VII.1).

Figure VII. 1 :

 1 Figure VII.1: Compilation of pre-eruptive and syn-eruptive diffusion timescales in ascending order modelled in several types of minerals in different volcanic systems, involving mafic or evolved magma compositions. Magma dynamics prior to these eruptions involve a magma recharge with or without magma mixing, they are shorter for basaltic or basaltic andesitic volcanoes (generally from months to a few years), while they are of several years to several decades for more evolved systems. Timescales modelled in several minerals from the same eruption are on the same abscissa.

Figure

  Figure VII.1 (continued): The eruptions studied in this thesis are highlighted in this compilation: Grand Fond (24 ka), Goodwill Entrance (18 ka), PPR1 (12.6 ka), PPR2 (12.5 ka), PPR3 (9 ka), for Dominica island and Bezymianny 1956 (blast and Plinian phase), Kizimen (2010-2013; for andesite "and" and dacite "dac"), in Kamchatka and for opx, mgt and plag. Other eruptions of the following volcanoes were combined here, in Alaska: Novarupta-Katmai 1912 (Singer et al., 2016), Shishaldin 1999 (Rasmussen et al., 2018); in the Andean Southern Volcanic Zone: Caburgua cones 5050 BC years (Morgado et al., 2017) and in Chile: Llaima 1850, 1957 and 2008 eruptions (de Maisonneuve et al., 2016; Ruth et al., 2018), Laguna del Maule Holocene 3,2 2,2 and < 1,8 ka (Andersen et al., 2018), Tatara-San Pedro complex Holocene eruptions and 235 ka (Costa and Chakraborty, 2004; Costa and Dungan, 2005), Taa´paca Volcano 1,5-0,5 to 0,45-recent (Rout et al., 2021); in the Canary Islands (Spain): El Hierro submarine eruption 2011-2012 (Longpre et al., 2014; Marti et al., 2013) and in Tenerife Siete Fuentes, Fasnia and Araf 1704-1705 (Albert et al., 2016, 2015); in Costa Rica: Irazú volcano 1723, 1963 and 1965 eruptions (Boyce and Hervig, 2009; Oeser et al., 2018; Ruprecht and Plank, 2013); Dominica with Layou eruptions (51 ka) and Roseau eruptions (33 ka) (Solaro et al., 2020); the East Pacific Rise 9°30'N and 10°30'N and 9°50'N 2005-2006 eruptions (Moore et al., 2014; Pan and Batiza, 2002); in Germany: Laacher See 12. Ka (Humphreys et al., 2021; Rout and Wörner, 2018; Sundermeyer et al., 2020b) and East and West Eifel 0.7-0.001 Ma (Sundermeyer et al., 2020b); in Greece: Santorini 25-39 ka, Santorini 21,8 ka (Druitt et al., 2012; Fabbro et al., 2018; Flaherty et al., 2018), Santorini 1600 BC Minoan or Late Bronze Age eruption (LBA) (Druitt et al., 2012; Flaherty et al., 2018); in Guadeloupe: La Soufrière 5680 BCE, 1657 CE, 1530 CE, 1010 eruptions and Échelle 341 CE (Metcalfe et al., 2021); Hawaiʻi: Mauna Loa 2,1 ka (Couperthwaite et al., 2020); Kīlauea 1500-1820 (Lynn et al., 2017); Kīlauea 1959 (Rae et al., 2016); Iceland: Borgarhraun 10.5-7 ka (Mutch et al., 2019), Bárðarbunga Middle and Early Holocene, 1477 AD (Caracciolo et al., 2021), Bardarbunga Pleistocene 10 ka (Mutch et al., 2021), Laki 1783-1784 (Hartley et al., 2016), Eyjafjallajökull 2010 (Pankhurst et al., 2018); in Indonesia: Younger Toba Tuff Eruption (Sumatra) 74 ka (Matthews et al., 2012), Kelud 1990 (Utami et al., 2021); Merapi 2006 and 2010 (Costa et al., 2013; Li and Costa, 2020); Italy: Etna 1669, 1991, 1993, 2006 and 2008 (Kahl et al., 2017, 2015, 2013, 2011), Vesuvius 79 CE, 1944 (Morgan et al., 2006, 2004), Stromboli eruptions < 2000 years (Petrone et al., 2018, 2016) and Campi Flegrei < 5 ka (Pelullo et al., 2022); Kamchatka : Bezymianny 2000, 2006, 2008, 2009, 2010, March and September 2012 (Davydova et al., 2018), Shiveluch 7600 BP (Gordeychik et al., 2018) ; La Réunion: Piton de la Fournaise 2007 (Albert et al., 2019) and 2014-2015 (Sundermeyer et al., 2020a); Mexico: Nevado del Toluca 26 ka (Weber et al., 2019); in New Zealand: Auckland Volcanic Field Late Pleistocene-Holocene (Brenna et al., 2018), Ruapehu 1969, 1971, 1977, 1995, 1996 (Kilgour et al., 2014), Taupo Volcanic Zone : Kidnappers and Rocky Hill (1 Ma ; Manganiko) (Cooper et al., 2017), Whakamaru 340 ka and Okataina Volcanic Center 50 ka (Matthews et al., 2012), Oruanui 25,4 ka and post-Oruanui rhyolites 2.15 ka (Barker et al., 2016) and Kaharoa 0,7 ka (Rubin et al., 2017); MORB (1982-1991 and 1993 flows of the Coaxial segment of the Juan de Fuca ridge, the 1996 flow of the North Gorda ridge, and from the Western Volcanic Zone of the ultra-slow spreading Gakkel Ridge) (Zellmer et al., 2011); USA: Long Valley 767 ka (Chamberlain et al., 2014), Yellowstone ca. 631 Lava Creek Tuff (Shamloo and Till, 2019), Yellowstone Scaup Lake 262 ka (Brugman et al., 2022; Till et al., 2015), Mount St Helens 1980 (Saunders et al., 2012) and Vanuatu: Ambae 2017 (Moussallam et al., 2019). Monitored eruptions are shown with an asterisk (*). The symbol position represents the pre-eruptive timescales given by the minerals (median of the dataset, when all data was available) and error bars correspond to the minimum and maximum timescale modelled. For some studies, not all the timescales modelled were provided in the articles (or only the timescales ranges were provided) so the values taken were the minimum and maximum timescales that were provided. Dashed lines highlight the longest timescale associated to a particular volcanic context (for example, "Last MORB", for the East Pacific Rise, around 130 days). The end of the graph shows timescales that only correspond to arc volcanoes and more or less evolved compositions.

  (Figure VII.1).

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FIGURE 1 |

 1 FIGURE 1 | Dominica and studied ignimbritic eruptions. (A) Lesser Antilles arc, with Dominica in the central part of the arc. (B) Dominica and location of the sampled outcrops (stars) for Layou and Roseau ignimbrites, after Boudon et al. (2017). Red: deposits from Morne Trois Pitons-Micotrin eruptive center, with dark red deposits corresponding to Roseau eruption. Blue: deposits from Morne Diablotins eruptive center, with dark blue deposits corresponding to Layou eruption. (C) Layou Plinian fallout deposit. (D) Roseau fallout deposit. (E) Layou pyroclastic density current deposit. (F) Roseau pyroclastic density current deposit.

  ). Whole rock samples were gently crushed to liberate crystals and sieved. Orthopyroxene (751 for Layou, 701 for Roseau) crystals were separated by hand-picking under binocular microscope and mounted in epoxy resin. Orthopyroxenes were oriented during mounting to obtain b-c sections needed for diffusion modeling along the b-axis of crystals. Frontiers in Earth Science | www.frontiersin.org November 2020 | Volume 8 | Article 574032 4

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Frontiers 7 FIGURE 3 |

 73 FIGURE 3 | Overview of opx textures and composition. Selected SEM images illustrating the opx zonations for Roseau (A: normal; B: reverse; C: multiple, reverse+normal) and Layou (D: normal; E: reverse; F: multiple, normal+reverse). Frequency histograms of the zoned and unzoned orthopyroxene compositions in Roseau (left side) and Layou (right side) expressed in terms of En (mol%). (G,J) unzoned crystals in grey; (H,K) cores of zoned crystals in blue for Layou and red for Roseau; (I, L) rims of zoned crystals, in dashed blue for Layou and dashed red for Roseau. Unzoned orthopyroxene compositions: single point analyses in the crystal core. Core and rim composition: averaged plateau compositions of the zoned crystals, with a minimum length of ∼15 μm. N: number of crystals.

FIGURE 4 |

 4 FIGURE 4 | Selected melt inclusion compositions as a function of host-orthopyroxene composition (En mol%). (A) SiO 2 ; (B) CaO; (C) FeO; (D) Na 2 O.

FIGURE 5 |

 5 FIGURE 5 | Examples of two modeled diffusion profiles for Roseau (A) and Layou (B). Red/blue line on each Back-Scattered Electron (BSE) image (left) indicates the profile position [Electron Probe Microanalyses (EPMA) full profile] on the respective orthopyroxene (Roseau, Layou). Diffusion profiles (right) are reported as Mg# [Mg/ (Mg + Fetot)], as a function of the distance from the center of the profile (in µm). Red rhombs: EPMA and BSE greyscale intercalibrated data; black line: diffusion model. Time (t) is given for each profile. The initial conditions are given by the dashed line (staircase profile).

FIGURE 6 |

 6 FIGURE 6 | Timescale estimate from Fe-Mg interdiffusion modelling in orthopyroxenes. Frequency histogram of Fe-Mg interdiffusion timescales modelled at 850 °C for the main core-rim boundary of (A) Roseau opx and (C) Layou opx; (B) and (D) individual ages and their absolute uncertainties based on the propagation of a temperature uncertainty of ±25 °Cf or(B) Roseau and (D) Layou. The inset graphs inside (B) and (D) represent selected individual ages from the peak in interdiffusion ages.

FIGURE 7 |

 7 FIGURE 7 | A variant Crystal Systems Analysis (CSA) diagram dedicated to opx, after the methodology of Kahl et al. (2015), demonstrating the direction and nature of sequential crystal zones of the opx from the Roseau (A) and Layou (B) eruptions. Each crystal is represented by a single arrow. Black arrows denote simple zoning and red arrows denote multiple zoning, from core to rim. The size of the magmatic environment boxes (ME1 to ME3, with subscript L for Layou and R for Roseau) are proportional to the number of unzoned crystals (grey) of the given composition (XEn) plus the number of zoned crystals (white) with cores of the given composition.

FIGURE 8 |

 8 FIGURE 8 | A dynamic magma model within the crust. Inset: Localization within the crust of the mush zone studied here. The color of the crystal is linked to magma temperature, from "cold" environment in blue to "hot" one in red, with intermediate temperature in orange; relative temperature estimates are on the basis of opx composition. Initial stage for (A) Roseau and (C) Layou. The dynamic model proposed on the basis of opx for (B) Roseau and (D) Layou.

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Figure

  Figure II.1: Subduction characteristics…

Figure II. 11 Figure

 11 FigureII.11

Figure II. 12 Figure

 12 FigureII.12

Figure II. 13 Figure

 13 FigureII.13

Figure II. 17

 17 FigureII.17

Figure II. 18 :

 18 FigureII.18: Seismic stations and Klyuchevskoy group.

Figure

  FigureII.18a

Figure II. 21 Figure

 21 FigureII.21

Figure

  FigureIII.18a 134

Figure III. 22 141

 22 FigureIII.22 141

II.10a1) with

  Kamchatka led to the extinction of the northern part of the former subduction zone and the formation of the current subduction zone 150 km to the east (Figure

II.10a2) (

	Lander and Shapiro,
	2007). It formed the first peninsula, the Shipunski segment (Figure II.10a2). A new subduction zone
	then formed 7 Ma ago and extended northwards for at least 5 Ma (Figure II.10a2). Thus, the activity
	of the old volcanic chain in western Kamchatka gradually died out and a new volcanic chain was formed
	in the east: the Eastern Volcanic Belt (Figure II.9a), equidistant from the old one (forming the
	Kronotski segment, Figure II.10a3). Near the Kamchatka-Aleutian junction, the subduction would
	have started 2 Ma ago (Kamchatski Mys joined Kamchatka, shaping the modern structure of the arc
	(Figure II.10a4). These events explain that the northern segment of Kamchatka has some features that
	are not present in the south (Kurile islands): a second Quaternary volcanic chain, whose activity ceased
	only in the Holocene and three peninsulas (Shipunski, Kronotski and Kamchatkaski Mys) located on
	the Pacific side of Kamchatka. Furthermore, the depth of the dipping panel decreases near this zone,
	from 650 km in the southern segment of the arc to 150-450 km beneath Kamchatka (dip 55° to 35°) and
	the subduction zone does not include deep earthquakes, which hinders understanding of the slab
	structure (Figure II.9b)

.9a; Figure II.11). They

	, Plosky Tolbachik,
	Bezymianny, hydrothermally active Ushkovsky, and other extinct ones (Portnyagin et al., 2007a)
	(Figure IIexhibit several types of compositions: either mainly basaltic
	(Klyuchevskoy, Ushkovsky, Tolbachik) or andesitic dacitic (Zimina) or purely andesitic (Bezymyanny,
	Udina)

III.2.4. Geochemical analyses III

  .2.4.1. Whole rock analyses Analyses of major element compositions of total rocks were performed by Inductively Coupled Plasma-Optical Emission Spectrometry (ICP-OES) at the Centre de Recherches Pétrographiques et Géochimiques (CRPG) in Nancy after acid attack of powders obtained in the laboratory(Carignan et al., 2001).

III.2.4.2.

Ponctual mineralogical analyses (crystal compositions and melt inclusions; EPMA)

Supplementary Table S4: Percentages of

			Fond	Goodwill	PPR1		PPR2		PPR3
	ME	ME1	ME2	ME1	ME2	ME1	ME2	ME1	ME2	ME1	ME2
	Unzoned (%)	59	26	32	8	33	2	46	3	30	6
	Population 1 SZ (%)	2		10		25		20		16
	Population 1 MZ (%) 2		14		18		16		26
	Population 2 SZ (%)		10		26		17		8		9
	Population 2 MZ (%)		1		11		5		7		13
	Total (%)	62	38	55	45	76	24	82	18	72	28

Eruption Crystal Fraction SiO 2 TiO 2 Al 2 0 3 FeO MnO MgO CaO Na 2 O K 2 O P 2 O 5 Total MI in opx

			(in µm)
	GE	L3C5_R	355	77,01 0,16 12,54 2,38 0,17 0,21 1,73 2,89 2,86 0,06
	GE	L2C6	355	75,53 0,30 12,87 3,13 0,10 0,22 2,06 2,69 3,09 0,01
	GE	L6C24_R	355	76,60 0,36 12,22 2,69 0,05 0,25 2,11 3,28 2,39 0,05
	PPR2	L3C22_R	355	75,32 0,26 11,84 3,83 0,10 1,09 2,34 2,53 2,22 0,48
	PPR2	L4C7	355	77,54 0,21 12,60 2,38 0,07 0,19 1,67 2,82 2,47 0,05
	PPR3	L9C1_R	355	77,25 0,15 12,32 2,09 0,11 0,16 1,60 3,23 3,09 0,01
	PPR3	L9C13_R	355	75,51 0,20 11,81 3,96 0,14 1,04 1,74 3,32 2,29 0,00
	PPR1	L7C30_c	250	77,05 0,21 12,31 1,87 0,14 0,26 1,88 3,42 2,83 0,02
	PPR1	L7C30_h2	250	51,74 0,15 0,60 26,85 0,97 18,48 1,09 0,04 0,09 0,00
	PPR1	L7C29mic	250	78,85 0,26 12,44 2,57 0,13 0,40 1,76 1,12 2,47 0,00
	PPR1	L7C29_h1	250	51,78 0,19 0,61 27,47 1,21 17,64 1,05 0,01 0,01 0,03
	PPR1	L7C29mi2c	250	77,16 0,14 12,30 2,40 0,13 0,21 1,73 2,94 3,00 0,00
	PPR1	L7C29_h2	250	51,98 0,19 0,61 26,86 1,16 18,13 1,03 0,05 0,00 0,00
	PPR2	L3C22mi_c	250	75,66 0,21 13,25 1,93 0,04 0,30 2,26 3,19 3,11 0,06
	PPR2	L3C22_h1	250	51,98 0,01 0,76 26,15 0,81 19,26 1,00 0,00 0,01 0,02
	PPR2	L4C11mi1_c	250	76,49 0,26 13,22 2,32 0,02 0,17 1,82 2,62 3,08 0,00
	PPR2	L4C11h1_h	250	51,15 0,08 0,48 27,47 0,99 18,75 1,07 0,00 0,00 0,00
	PPR2	L5C19mi_c	250	79,26 0,23 11,35 2,27 0,07 0,15 1,46 3,00 2,15 0,07
	PPR3	L9C18mi_c	250	77,04 0,25 12,21 2,51 0,09 0,19 1,53 3,38 2,81 0,00
	PPR3	L9C18_h	250	51,20 0,10 0,50 27,89 1,09 18,19 0,94 0,04 0,03 0,02
	PPR3	L13C2mi1_c	250	75,70 0,00 12,76 3,48 0,03 0,21 1,88 3,10 2,87 0,00
	PPR3	L13C2_h	250	51,26 0,14 0,43 29,00 1,03 17,08 1,01 0,02 0,04 0,00

4; Supplementary Data V.3).

Allan et al. (2013)

fO2 correction), following the fO2 dependence in equation (V.4). The diffusion coefficient of

Dohmen et al. (2016)

was also tested, using the open-access fitting algorithm for zoned opx of

[START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF]

, even if this algorithm is only based on grayscale values and not on EPMA data. This algorithm is the same as the one originally developed by

Morgan et al. (2004)

and modified by

Petrone et al. (2016)

to resolve multiple diffusive boundaries.

The model used to estimate the timescales with

Ganguly and Tazzoli (1994)

diffusion coefficient is the equivalent of AUTODIFF

(Allan et al., 2013; Couperthwaite et al., 2020; Hartley et al., 2016; Ostorero et al., 2021, accepted; Pankhurst et al., 2018)

but for opx (see the demo version for olivines in

Couperthwaite et al. (2020)

and Supplementary Note for more details). Profiles in which Al2O3 gradients are decoupled from the Mg number (Mg# = Mg Mg+Fe tot ), are modelled so that the most of the gradients are not due to growth of the crystals but diffusion of Fe-Mg

(Flaherty et al., 2018)

(Supplementary Figure V.

.1; Supplementary Figure V.17).

  For the Plinian phase, they are 40 times longer than D ofGanguly and Tazzoli (1994) with fO2 (Appendix TableV.

1; Supplementary Figure V.17).

			Blast opx			Plinian phase opx		Blast mgt	Plinian phase mgt
			D (G&T)				D (G&T)
		T mgt (S)	T mgt (G&E)	With fO2 (T mgt (S))	D (D) with fO2	T mgt (S)	T mgt (G&E)	T opx (P)	With fO2 (T mgt (S))	D (D) with fO2	T mgt (S)	T mgt (S)
	Median
	timescale	51.6	29.5	4.8 28.5 12.8	7.2	47.1 1.5 57.8 44.6	2.8
	(days)
	Expected
	timescales	83.7	45.3	9.2 62.7 25.8 14.4	87.7 2.4 93.1 59.7	3.6
	(days)
	Expected
	timescale std. Error	0.3	0.3	0.3	0.3	0.3	0.3	0.3	0.3	0.3	0.7	0.4
	(days)

Appendix Table V.1: Summary of the

  1* , Hélène Balcone-Boissard 2 , Georges Boudon 1 , Nikolai M.Shapiro 3,4 , Alexander Belousov 5 , Marina Belousova 5 , Andreas Auer 6 , Sergey L. Senyukov 7 , Svetlana Ya. Droznina 7 1 Université Paris Cité, Institut de physique du globe de Paris, CNRS, F-75005 Paris, France 2 Institut des Sciences de la Terre de Paris (ISTeP), UMR 7193, CNRS-Sorbonne Université, Paris, France 3 Institut des Sciences de la Terre (ISTerre), Université Grenoble Alpes, CNRS, 38400, Grenoble, France 4 Schmidt Institute of Physics of the Earth, Russian Academy of Sciences, 10 B. Gruzinskaya St.,

	Moscow, 123995
	5 Institute of Volcanology and Seismology, 9 Piip Boulevard, Petropavlovsk-Kamchatsky, 683006,
	Russia

6

Department of Geoscience, Shimane University, 1060 Nishikawatsu, Matsue 690-8504, Japan

7

Kamchatka Branch of the Geophysical Survey, Russian Academy of Sciences,

9 Piip Boulevard, 683006, Russia

  . However, plagioclase rims in the andesitic bands of banded samples of the same eruption are of different types but calcium-rich rims are mainly found (reverse-zoned). Plagioclase in the andesites then need to be studied, to see if normal-zoned plagioclase crystals are dominant, alongside the reversely-zoned opx. The

	mean variation of ∆ 𝐾 𝐷 𝐹𝑒 2+ -𝑀𝑔	𝐹𝑒 2+ -𝑀𝑔 𝐾 𝐷	from the cores to the reverse-zoned rims of the opx,

). Image from the ArticDEM: Hillshade Gray from the ArcGIS Living Atlas of the

	World	(by	Esri	Imagery,	courtesy	of	the	Polar	Geospatial	Center;

https://livingatlas2.arcgis.com/arcticdemexplorer/) 117 and modified according to other studies

118,119 . b)

Sampling sites on Kizimen (Supplementary Table 1) (Image credit: Nasa Earth Observatory image from June 17, 2014, using Landsat 8 data from the United States Geological Survey (USGS) Earth Explorer; courtesy of Robert Simmon). c) Dacite (64 wt% SiO2; 5.6 wt% Na2O +K2O). d) Andesite (58-60 wt% SiO2; 4.8-5 wt% Na2O +K2O). e-f) Banded samples

2.6b-7b; Table VI.2.2; Supplementary Data VI.2.2).

2.2).

Mgt in the andesite mainly display a reverse zoning in Ti (Supplementary Data VI.2.2). Mgt in the dacite and andesite appear to have similar timescales, with short timescales (< 15 days; 14 timescales) and even shorter ones for the dacite (less than 5 days; 3 timescales) (Figure VI.

Table VI .

 VI

	Crystal K8 250L4C4	Timescales (days) 1825	Zonation type NZ+RC	Error + (days) 730	Error -(days) 7300	An change An78-84 (intermediate zone) to An62-79
	K8 250 L2C6 K5 500_L1C2	10 1825	RZ NZ+RC	2 365	40 5475	An42-52 (intermediate zone) to An60-70 An78-84 (intermediate zone) to An62-80
	K5 250_L7C21 K8 250 L5C24	20 1825	RZ NZ+RC	4 365	100 5475	An42-52 (intermediate zone) to An60-71 An78-84 (intermediate zone) to An62-81
	K8 250L6C20	30	RZ	5	100	An42-52 (intermediate zone) to An60-72
	K8 250L9C5	30	RZ	5	200	An42-52 (intermediate zone) to An60-73
	K5 250 L1C4	30	RZ	5	100	An42-52 (intermediate zone) to An60-74
	K8 250L3C11	40	RZ	4	100	An42-52 (intermediate zone) to An60-75
	K8 250L9C13	40	RZ	5	150	An42-52 (intermediate zone) to An60-76
	K8 250 L8C12A	40	RZ	5	150	An42-52 (intermediate zone) to An60-77
	K5 250 L5C3	40	RZ	5	200	An42-52 (intermediate zone) to An60-78
	K8 250 L10C15	50	RZ	10	200	An42-52 (intermediate zone) to An60-79
	K5 250_L2C12	175	RZ	30	710	An42-52 (intermediate zone) to An60-80
	K8 250 L2C7	250	NZ+RC	75	1460	An78-84 (intermediate zone) to An62-70
	K8 250 L8C12B	365	RZ+RZ			An42-52 (intermediate zone) to An60-80
	K5 250_L2C5	730	NZ+RC	200	3650	An78-84 (intermediate zone) to An62-70
	K5 500_L3C4	730	NZ+RC	200	3650	An78-84 (intermediate zone) to An62-71
	K8 250L9C6	730	NZ+RC	200	1025	An78-84 (intermediate zone) to An62-72
	K8 250L9C14	730	NZ+RC	200	1825	An78-84 (intermediate zone) to An62-73
	K8 259 L5C8	730	NZ+RC	200	2920	An78-84 (intermediate zone) to An62-74
	K8 250 L10C22	730	NZ+RC	200	2920	An78-84 (intermediate zone) to An62-75
	K8 250 L12C8	730	NZ+RC	200	3650	An78-84 (intermediate zone) to An62-76
	K5 250 L2C16	730	NZ+RC	150	2920	An78-84 (intermediate zone) to An62-77
	K8 250L9C31	1095	NZ+RC	200	2555	An78-84 (intermediate zone) to An62-78
	K8 250 L4C17	1095	NZ+RC	365	5475	An78-84 (intermediate zone) to An62-78
	K8 250L8C11	1325	NZ+RC	365	7300	An78-84 (intermediate zone) to An62-78
	K8 250L2C13 (1)	1460	NZ+RC	365	3650	An78-84 (intermediate zone) to An62-78

(Figure VI.2.8).

Table VI .

 VI we decided to perform the experiment at a temperature below the liquidus of the feldspar and spinel for the synthetic basalt (around 1227 and 1258°C respectively, at NNO + II conditions) and above them for the dacite (around 1150 and 1153°C respectively, at NNO + II conditions)[START_REF] Gualda | Rhyolite-MELTS: a Modified Calibration of MELTS Optimized for Silica-rich, Fluid-bearing Magmatic Systems[END_REF]. The experiment was then run at a temperature of 1200°C, at ambient pressure. At 1200°C, a crystal fraction of 20 vol% in the synthetic basalt is expected (with 13 vol% of feldspar) (Supplementary FigureVI.3.2) [START_REF] Gualda | Rhyolite-MELTS: a Modified Calibration of MELTS Optimized for Silica-rich, Fluid-bearing Magmatic Systems[END_REF].

	(Table VI.3.1; Supplementary Note). Powders from each of the two end-members were melted and homogenized for 4 h inside a high-temperature furnace (Nabertherm® HT 04/17) at ambient pressure at the Department of Physics and Geology (University of Perugia) at 1500 °C in Pt80Rh20 crucibles (following the method of Morgavi et al. (2013a, b, 2015)). After quenching the samples in air, homogeneous bubble-and crystal-free glasses were obtained and used as starting materials for the mixing experiment (Supplementary Figure VI.3.1). Based on MELTS calculations of the crystal fraction for both end-members (Supplementary Figure KIZ-96-01/1 (wt%) Synthetic basalt Error min Error max Dacite Error min Error max SiO2 49.7 51.28 0.66 0.85 64.16 0.67 0.69 TiO2 1.22 1.25 0.34 0.28 0.61 0.19 0.13 Al2O3 18.84 19.25 0.64 0.71 16.58 0.57 0.41 Fe2O3 5.32 FeO 5.56 10.55 0.63 0.54 5.26 0.59 0.35 MnO 0.19 0.02 0.08 0.08 0.11 0.15 0.14 MgO 5.2 5.32 0.31 0.31 2.38 0.12 0.16 CaO 9.25 8.43 0.39 0.53 5.34 0.2 0.28 Na2O 2.74 3.08 0.24 0.41 3.64 0.54 0.25 K2O 0.76 0.76 0.13 0.18 1.76 0.21 0.28 P2O5 0.17 0.04 0.08 0.14 0.17 0.19 0.15 LOI 0.73 VI.3.2), Oxide Total 99.67 100 100

3.5b-d; Supplementary Figure VI.3.3; Supplementary Table VI.3.2).

  .3.1), with an acicular dominant shape of crystals for the andesitic band and tabular for the dacite and a mean orientation of the crystals around 72° (Figure VI.The crystal number density 𝑁 𝑎 versus the crystal microlite content 𝜑 show that both growth and nucleation were abundant for the andesite whereas the dacite experienced an abundant nucleation without growth (Supplementary FigureVI.3.3). CSD curves evidence that the crystals studied in the natural samples are small, with crystals ranging from 4.10 -4 to 0.01 mm for the two bands (characteristic lengths of crystals of 0.9 µm and 0.4 µm, respectively) (Figure VI.

3.6a-b; Supplementary Table VI.3.2).

  The CSD curves are concave upwards(Figure VI.3.6b), which require two linear regressions to better approximate the general trends of the curves. It implies that at least two crystal populations nucleated and grew at different regimes in the natural samples, with one being in a growth-dominated regime and another in a nucleation-dominated regime

3.6b-d; Table VI.3.3).

  for the two different parts of the CSD curve (Figure VI.For the dacitic band, using the maximum and minimum growth rates of Hammer and Rutherford (2002), longer residence timescales are found and are of 41 ± 144 h and of 30 ± 107 days, for the two different parts of the curve (Figure VI.

3.6c-d; Table VI.3.3).

  Timescales are different depending on the experimentally growth rates used. As the melt in the andesitic and dacitic magmas was rhyolitic (residual glasses analyses, FigureVI.3.4), the same growth rate determined for a rhyolitic melt seem to be more appropriate. Very short timescales from minutes to 2 hours were obtained for the two parts of the andesite and dacite (13 ± 9 min and 48 ± 32 min and of 6 ± 4 min and 2 ± 1 h, respectively)[START_REF] Mollard | Decompression-induced Crystallization in Hydrated Silica-rich Melts: Empirical Models of Experimental Plagioclase Nucleation and Growth Kinetics[END_REF] (Figure VI.

3.6c-d; Table VI.3.3).

  So, a nucleation event occurred in the order of hours for both the andesite and dacite, followed by a growth-dominated regime, in the order of several minutes.

	Sample	G (mm/h) ; Higgins et al. (2000)	GL (mm/h) ; Agostini et al. (2013)	Nuclea tion rate J (mm -3 /h)	Nuclea tion rate J (mm -3 /h)	τ (min/max (average ± std dev) -using the two parts of the CSD curve)	τ (min/max -using the two parts of the CSD curve)
	Experime nt	1.24.10 -2 2.75.10 -2	0.25	0.11	-	-
	Natural banded lava -Andesitic band	-	-	-	-	2-22/9-81 min (12 ± 14/45 ± 51 min) (Shea and Hammer, 2013)	7-19/26-70 min (13 ± 9/48 ± 32 min) (Mollard et al., 2012)
	Natural					4 min -81 h/73 min -60	3-9 min /56 min -
	banded lava -	-	-	-	-	days (41 ± 144 h/30 ± 107 days) (Hammer and	3 h (6 ± 4 min/2 ± 1 h) (Mollard et
	Dacite					Rutherford, 2002)	al., 2012)

Table VI .

 VI

Table VI .3.1. Timescales are reported in Table VI.3.3.

 VI

; Supplementary Figure VI.3.3; Supplementary Table VI.3.1-2).

  These crystals show an abundant growth but limited nucleation during the experiment, with the lowest 𝑁 𝑎 , of ~140 mm -2 , compared to the natural banded lavas (Supplementary FigureVI.3.3). The crystals in the experiment then did not nucleate much but had time to grow. The CSD curve shows a size distribution of the crystals from 0.02 mm to 1.2 mm, with a characteristic length of 0.15 mm, which are bigger than the ones investigated in the natural banded lavas (Figure VI.

3.6a; Supplementary Table VI.3.2).

  The CSD curve of the experimental sample can be approximated as a straight line (Figure VI.3.6a), it implies that nucleation and growth were continuous during crystallization. Compared to Supplementary Figure VI.

3.A.2 Geochemical compositions of the glasses, transects and crystals VI

  .3.A.2.1. Whole rock and glass compositions These transects were done to cover different areas of the mixing patterns developed in the experiment: crossing different mafic filaments with isolated crystals in them or without crystals and also in felsic enclaves (Figure VI.

	In the Total alkali diagram, the starting materials compositions are in the basaltic and dacitic field,
	respectively (51.3 wt% SiO2 and 4 wt% Na2O+K2O and 64 wt% SiO2 and 5.4 wt% Na2O+K2O) (Figure
	VI.3.4).
	The main glasses of all zones in the experimental sample are between 51-66 wt% in SiO2 and 3.5-6 wt%
	in Na2O+ K2O (Figure VI.3.4).
	VI.3.A.2.2. Transects

1 𝑠𝑙𝑜𝑝𝑒

).

VI.

The compositional variability was studied along several compositional transects (Supplementary Table

VI.3.3

).

3.8a1-d3, Supplementary Figure VI.3.5). Figure VI.3.8 shows

  oscillatory patterns as passing through filaments of the two different melts, with lower compositions in SiO2 than the dacite but higher than the synthetic basalt. The mafic compositions are not the same depending on the three transects (for example, on Figure VI.3.8a1-a3, SiO2 content is mainly varying from 60, 57, 59 wt%, respectively and 3.6, 5 and 3.5 wt% in MgO) whereas the felsic components are between 64-65 wt% for SiO2 respectively and around 2.5 wt% in shows the compositions of the transects, with the SiO2 content mainly varying from 62-65 wt%, CaO content between 5-6 wt%, MgO between 2-3 wt% and FeO between 5-6 wt%. Some variations occur, with lower values reached for the SiO2, MgO and FeOtot content and higher values for the CaO content Practically no mafic glass is present in the range of the compositions from the synthetic basalt to the most mafic glass generated by the experiment, with only a few points being in this interval (from 51-55 wt% in SiO2, for example).

	(Figure VI.3.8a4-d4). Mafic compositions have a very low representativity as seen on Figure
	VI.3.8a4-d4, with only a few points between 51-53 wt% in SiO2 and between 5-6 wt% in MgO.
	Harker diagrams show the compositional variability of these glasses, with the compositions mainly
	covering the areas between 56-65 wt% in SiO2, 5-8 wt% in CaO, 2-6 wt% in MgO, 1-2.5 wt% in K2O,
	2.5-4 wt% in Na2O and 15-18 wt% in Al2O3 (Figure VI.3.9).

,

  Wo content (𝑊𝑜 = 𝐶𝑎 𝐶𝑎+𝑀𝑔+𝐹𝑒 ) between 33-45 % and 2 crystals with Wo 6-7 %.

	Sample	Number of crystals	ϕ	Na (mm -2 )	Best-fit crystal habit (x, y, z)	score (R 2 )	S/M	S/L M/L
	Experiment	458	0,23	138	1 : 4 : 9	0,88	0,25	0,11 0,44
	Natural
	banded lava -Andesitic	490	0,27	26382	1 : 1.4 : 2.8	0,82	0,71	0,36 0,50
	band
	Natural
	banded lava -	441	0,02	31148	1 : 2.5 : 5	0,81	0,40	0,20 0,50
	Dacite
	Supplementary

Table VI .3.1: Number of crystals, microlite crystal fraction 𝝋, two-dimensional crystal number density 𝑵 𝒂 , best-fit crystal habit using CSDSlice (Morgan and Jerram, 2006) for the experimental and natural samples and the associated R 2 and the aspect ratios (short axis over medium axis (S/M), short axis over long axis (S/L) and medium axis over long axis (M/L)).

 VI This summary presents the first results and interpretations of the second chaotic mixing experiment performed at the Department of Physics and Geology (University of Perugia). This study is in collaboration with Daniele Morgavi, Fabio Arzilli, Hélène Balcone-Boissard, Georges Boudon, Maurizio Petrelli, Tatiana Churikova and BorisGordeychik.

	Sample	Mean Orientatio n (°)	Ratio orientat ion ellipse axes	Alignm ent factor	Nv (mm -3 )	Lc (mm) (1σ error)	Inter ) cept (1σ error	Slope (1σ error)
	Experiment	-58	0,75	0,28	1,10.10 3	0,1505 (0,02)	9,25 (0,08)	-7.69 (0.25)
								Part 1	Part 2
	Natural
	banded lava -Andesitic	72	0,92	0,09	3,12. 10 7	0,0009 (0)	24,59 (0,07)	-1824 -493
	band
	Natural banded lava -Dacite	71	0,8	0,22	8,65.10 7	0,0004 (0)	27,96 (0,13)	-4020 -224

Table VI .4.1: Summary of the experimental conditions and the temperatures of the first melting process (first at 1500 °C), and then at the temperature where a crystal fraction of 13 vol% of feldspar may be generated at hypothesized NNO + II oxidizing conditions (simulated with MELTS) (Gualda et al., 2012). Liquidus temperature for feldspar as well as the amount of undercooling are specified.

 VI

			Liquidus for fd at ambient P (°C)	Undercooling compared to Liquidus (°C)	Oxidizing conditions	Time in the furnace before experiment
	Basaltic andesite	1172 13	1203	31	NNO + II 7h
	Dacite	1115 13	1150	35	NNO + II 7h 20 min
	The apparatus used for magma mixing experiments is the Chaotic Magma Mixing Apparatus (hereafter
	COMMA), same as the one presented in Chapter VI.3 and in Morgavi et al. (2015). The proportion of

felsic end-member was higher than the mafic one (80 vol % vs 20 vol %) (Table VI.

4.2)

. At the end of the experiment, rapid quenching of the experimental melt to form a glassy experimental product is performed (cooling rate of 41 °C/min)

[START_REF] Rossi | Exponential decay of concentration variance during magma mixing: Robustness of a volcanic chronometer and implications for the homogenization of chemical heterogeneities in magmatic systems[END_REF]

. After the experiment, quenched glasses are cored out from the outer cylinder and set in epoxy to obtain slices of 2-4 mm thickness. The patterns

59 (n = 63) 2 yrs (n = 37)

											Pressu
	Localizat ion	Eruption	Style	VE I	Volu me of ejecta (km 3 )	Av. Bulk composit ion SiO2 (wt%)	Av. MI (wt%)	Av. RG (wt%)	Temperat ure T (°C)	T determinat ion (°C)	re (MPa) ; depths	En conte nt opx	Median timesca les	Magmatic processes	Correlatio n with monitorin g data
											(km)
		Grand Fond	Ignimbr ite	5	4-5 a	62.9 a	78.1 b	77.5 b	850 ± 25 c Experiment al study c	300-400 b,c ; 12-16	En46-59 49) (n =	34 yrs (n = 22)	Magma mixing
	Dominica	Goodwill PPR1 PPR2 PPR3	Plinian	4	0.3-0.5 a	62.7 a 62.9 a 63.4 a 63.3 a	76.4 d 76.4 d 76.7 d 76.6 d	75.8 d 76.2 d 76.3 d 76.4 d	895 ± 9 (n = 5) d 893 ± 9 (n = 9) d 893 ± 10 (n = 12) d 886 ± 6 (n = 4) d	Melt inclusions e Melt inclusions e Melt inclusions e inclusions e Melt	50-200 d ; 2-8	En46-59 (n = 72) En46-59 (n = 59) En46-59 (n = 61) En46-	6 yrs (n = 35) Vigorous magma mixing in 2 yrs 4 yrs (n = 9) prior to the eruption in a 4 yrs (n = 28) temperatu re-zoned reservoir	Unmonito red
	Kamchatk a	Kizimen 2010	Extrusi ve-explosi ve	3 f	0.4	Dacite: 64.2 Andesite: 58.7	Dacite: 80.9 Andesi te: 78.4	Dacite and Andesi te: 77.7	Dacite: 832 ± 37 (n = 3) Andesite: 847 ± 57 (n = 21)	Mgt/ilm pairs g	123-290 h ; 6-11	En49-73 (n = 142 for the andesi te and 78 for the dacite )	Dacite: Opx -154 days (n = 16) Mgt -3 days (n e: Opx Andesit = 4)	Magma mixing in the andesite with degassing and the melt water in nt in enrichme	Yes, mainly with seismic data
										2015)
	(Figure VII.1).

Table VII .1: Summary of the characteristics for each eruption and results obtained in this thesis (in bold letters; Ostorero et al. (2021, accepted) and Chapter V). Compositions of whole rocks, melt inclusions, residual glasses of the volcanic samples studied are provided, from previous studies or from our results, as well as temperature and pressure. En content of the orthopyroxenes and mean timescales found with opx or magnetite are shown, as well as their possible correlation with monitoring data, when available. For Kizimen, the VEI was estimated to be of 3 (0.4 km 3 of ejecta, eruptive column of 10 km (Global Volcanism Program, 2013; Newhall and Rast, 1970; Senyukov et al., 2011)). For Bezymianny, the VEI was estimated to be of 5 (~2 km 3 of ejecta, eruptive column of 35-40 km and cataclysmic explosion with a laterally directed blast (called "blast" in the Table VII.1) (Global Volcanism Program, 2022; Newhall and Rast, 1970)). Average melt inclusions are given in opx-hosted melt inclusions. a Boudon et al. (2017); b Balcone- Boissard et al. (2018); c Solaro et al. (2019); d d'Augustin et al. (2020); e Putirka (2008); f Global Volcanism Program (2013); g Sauerzapf et al. (2008); h Auer et al. (2018); i Global Volcanism Program (2022); j Ghiorso and Evans (2008); k Yavuz (2021); l Newmann and Lowenstern (2002); m Turner et al. (2013).

 VII

												-119	in the
												days	dacite 1.5
												(n = 53) 22 days Mgt -	years before.
												(n = 11) Plag -730
												days (n
												= 28)
	Bezymia nny 1956	Vulcani an eruption s and in March 1956: laterally directed blast followe d by a Plinian column	5 i	2	Blast: 59.9 Plinian phase: 59.6	Blast: 75.8 Plinian phase: 76.3	Blast: 73.5 Plinian phase: 75.6	Blast: 915 ± 17 or 935 ± 19 (n = 6) Plinian phase: 991 ± 44 or 1020 ± 50 (n = 20) Plinian phase MI opx: 928 ± 26 (n = 20)	Blast and Plinian phase: mgt/ilm pairs g,j,k MI e	Plinia n phase: 200-400 l ; 9-15. Deepe r one inferre d from anothe r study 19-26 m )	En58-68 (n = 59 for the blast and 90 for Plinia n phase )	Blast: Opx -64 days (n = 30) Mgt -46 days (n = 10) Plinian phase: Opx -12 days (n = 39) Mgt -4 days (n = 29)	Magma before. convectio n 6 months-1 year before the eruption, cryptodo me storage and degassing or heating by magma injection 2 -3 months	Not clear as only one remote seismic station (pre-eruptive seismic crisis related to the Vulcanian activity).

TABLE 1 |

 1 Statistic repartition of zonations in opx from Layou and Roseau eruption. Opx are separated in the 315-500 µm fraction.

	LAYOU		751 opx
	Unzoned (%)	85
	Zoned (%)	15
		Reverse (%)	44
		Normal (%)	28
		Multiple (%)	28
			Normal + reverse (%)	76
			Reverse + normal (%)	24
	ROSEAU		701 opx
	Unzoned (%)	85
	Zoned (%)	15
		Reverse (%)	69
		Normal (%)	8
		Multiple (%)	23
			Normal + reverse (%)	12
			Reverse + normal (%)	88

TABLE 3 |

 3 Comparison of diffusion timescales obtained for different large volume silicic eruptions.

	Frontiers in Earth Science | www.frontiersin.org								Solaro et al.
		Location	Volcano	Eruption (age)	Magma	Dated process	Method	Diffusion timescales	Reference
					volume (km 3			(years)
					DRE)
		Dominica	Morne diablotins	Layou	∼5	Reservoir heating/magma mixing	Fe-Mg in opx	3-10	This study
		Dominica	Morne trois pitons-	Roseau	∼5	Reservoir heating/magma mixing	Fe-Mg in opx	3-10	This study
			micotrin
		Greece	Santorini	Minoan (1600 BC)	80	Magma recharge	Mg in plg	10-100	(Druitt et al., 2012)
		New zeland Taupo volcanic zone	Whakamaru (340 ka)	1,000	Magma recharge	Ti in quartz	∼10-60	Matthews et al., 2012
	14	New zeland Okataina caldera complex	Earthquake flat (50ka)	10	Magma recharge	Ti in quartz	∼100	Matthews et al., 2012
		New zeland Taupo volcanic zone	SG3 (<2.15 ka)	∼105 (bulk)	Reservoir heating	Fe-Mg in opx	<120	(Barker et al., 2016)
		New zeland Taupo volcanic zone	Oruanui (25.4 ka)	530	Melt extraction	Fe-Mg in opx	∼230	(Allan et al., 2013)
		Unites	Long valley caldera	Bishop tuff (767 ka)	600	Reservoir rejuvenation	Fe-Mg in opx	<500	(Chamberlain et al.,
		States							2014)
			--	-	-	Ti in quartz	--
			--	-	-	Ba and Sr in sanidine	--
		USA	Yellowstone	Central plateau Member (170-60	40-70	Reservoir rejuvenation	Zr, sanidine	<1kyrs	Stelten et al., 2015
				ka)
		Greece	Santorini	Minoan eruption, late bronze age	30-80	Repetitive melt injection, cooling,	Mg-Fe diffusion in opx,	<2 (last event)	(Flaherty et al. (2018))
						mixing	cpx
	November 2020 | Volume 8 | Article 574032								Dynamics Magma Plumbing System Dominica

  Le Chapitre IV et l'Annexe A sont publiés avec la Licence Creative Commons licence Attribution 4.0 International (CC BY 4.0) : http://creativecommons.org/licenses/by/4.0/. Avec cette licence, nous pouvons partager et adapter le matériel utilisé pour n'importe quel usage, même commercial, si les sources sont citées (ce qui est fait ici dans la thèse). Le Chapitre VI.1 est accepté dans la revue Communications Earth & Environment et a été publié le 19 novembre 2022 ; le pre-print sur Research Square est aussi publié avec la licence CC BY 4.0.

Ces échantillons ont ensuite été préparés selon un protocole bien défini, commençant par le nettoyage et le concassage des roches, afin de les broyer finement (analyses de roche totale) ou de les trier sous loupe binoculaire pour séparer les différents types de cristaux en fonction de leur taille. Pour ma thèse, l'étude étant principalement basée sur les chronomètres des cristaux d'orthopyroxènes et de magnétites, ceux-ci ont donc principalement été montés sur des plots et imprégnés dans de la résine. Les cristaux ont été observés au microscope électronique à balayage (MEB) afin de mettre en évidence l'éventuelle présence de zonations visibles (bordures de compositions différentes du coeur du cristal), témoins de changements dans les conditions magmatiques de stockage des cristaux avant une éruption. Leurs compositions sont étudiées grâce à des analyses à la microsonde électronique, sous la forme de profils allant de la bordure du cristal jusqu'à son coeur, dans les orthopyroxènes et dans les oxydes de fer et titane. Des analyses ponctuelles ont aussi été réalisées dans les coeurs des cristaux non zonés d'orthopyroxènes ou des oxydes.Ensuite, une description théorique des processus de diffusion et de ses principes sont présentés : description de la diffusion du point de vue microscopique avec les mécanismes d'échanges dans un cristal puis description de la diffusion du point de vue macroscopique, en abordant la loi de Fick qui permet de décrire le processus de diffusion dans l'espace et le temps ainsi que les solutions principales des équations de la diffusion pour un solide(Costa and Morgan, 2010;[START_REF] Streck | Mineral Textures and Zoning as Evidence for Open System Processes[END_REF]. L'étude des temps de diffusion des éléments dans les minéraux est une méthode actuellement utilisée pour obtenir des contraintes temporelles sur la dynamique dans les réservoirs magmatiques. Différents paramètres influencent la diffusion, comme la température, la pression, la fugacité d'oxygène et de l'eau, la composition du cristal et son anisotropie)[START_REF] Anderson | Probable relations between plagioclase zoning and magma dynamics, Fuego Volcano, Guatemala[END_REF][START_REF] Giacomoni | Plagioclase as archive of magma ascent dynamics on "open conduit" volcanoes: The 2001-2006 eruptive period at[END_REF] Petrone and Mangler, 2021;[START_REF] Streck | Mineral Textures and Zoning as Evidence for Open System Processes[END_REF]. Le degré d'influence de chaque paramètre sur la diffusion est discuté. L'étude de la diffusion dans les cristaux d'orthopyroxènes, abondants dans les systèmes magmatiques, est ensuite présentée, avec une description détaillée des coefficients de diffusion pour ces cristaux, ainsi que le modèle de diffusion utilisé dans cette étude pour modéliser l'interdiffusion Fe-Mg dans les orthopyroxènes. Ces cristaux sont à la base des différentes études sur le volcan Morne Trois Pitons-Micotrin de l'île de la Dominique, mais aussi du Bezymianny et Kizimen auKamchatka. La diffusion

Enfin, la méthode de la partie expérimentale de cette thèse est présentée. Celle-ci a consisté en la réalisation d'expériences de mélange de magmas à haute température à l'Université de Pérouse (Italie) au sein du Département de Physique et Géologie. Les expériences consistent à faire fondre deux verres de compositions différentes puis à les mélanger à haute température dans un creuset avec un système de rotation de deux cylindres dans des directions opposées, ce qui permet de déclencher un mélange chaotique(Morgavi et al., 2015). L'analyse des échantillons expérimentaux est ensuite présentée, en utilisant principalement le MEB et la microsonde électronique..

Scientific Reports | (2021) 11:11440 | https://doi.org/10.1038/s41598-021-90831-1

Frontiers in Earth Science | www.frontiersin.org November 2020 | Volume 8 | Article 574032

Acknowledgements

We would like to thank S. Hidalgo for her help with sample preparation, M. Fialin and N. Rividi for assistance during EPMA analyses and O. Boudouma for SEM imaging. We also would like to thank F. Costa and E. Kaminski for discussion on magma dynamics. This project was funded by the Institut de physique du globe de Paris (Doctoral grant from the French Ministère de l'Enseignement Supérieur de la Recherche et l'Innovation) and ANR V-Care-18-CE03-0010. Some data on Grand Fond eruption are part of the PhD thesis of C. Solaro-Müller 44 . We would like to thank the editor L. Pappalardo and reviewer G. Fabbro for their constructive and helpful comments.

Acknowledgments

We thank the editors E. Liu and J. Aslin, and T. Ubide as well as an anonymous reviewer for their constructive comments and suggestions that helped to improve the manuscript. We would like to thank S. Hidalgo for her help with sample preparation, M. Fialin and N. Rividi for assistance during EPMA analyses, as well as O. Boudouma and S. Borensztajn for SEM imaging and E. Delairis for the thin sections. We also thank D. J. Morgan for providing the files to model the timescales of the opx. We also acknowledge that the image from Figure 1a is from a Digital Elevation Model provided by the Polar Geospatial Center under NSF-OPP awards 1043681, 1559691, and 1542736 and Figure 1b from NASA Earth Observatory images by R. Simmon, using Landsat 8 data from the USGS Earth Explorer. Thank you also to R. Bajou for the discussion on the cumulative number of earthquakes and to F. Costa and S. B. Utami for the discussions on KD. This project was funded by the Institut de physique du globe de Paris (Doctoral grant from the French Ministry of Higher Education and Research and Innovation) and ANR V-Care-18-CE03-0010. The work of N. Shapiro, S. Senyukov, and S. Droznina has been supported by the Ministry of Science and Education of the Russian Federation under project no. 14.W03.31.0033 "Geophysical research, monitoring and forecasting of catastrophic geodynamic processes in the Russian Far East". The work of N. Shapiro has been also supported by the European Research Council under the European Union Horizon 2020 research and innovation program (Grant Agreement 787399-SEISMAZE). An open access license has been selected. the data and the writing of the manuscript text. N.S., S.S. and S.D. analyzed the seismic catalog of Kizimen and participated in the interpretation of the seismic data. All authors contributed to the writing of the manuscript text.

VI.3.5. Acknowledgments

L. Ostorero would like to thank the University Paris Cité and Institut de physique du globe de Paris for the international mobility grants that enabled this work to take place and D. Perugini and the members of the department of Physics and Geology of the University of Perugia for two stays during her PhD. We also thank S. Hidalgo and E. Delairis for their help with sample preparation, O. Boudouma, S. Borensztajn and A. Di Michele for SEM imaging as well as M. Fialin and N. Rividi for their help during EPMA analyses. O. Bernard is also thanked for his help and the several discussions on the CSD method. We also acknowledge that the image from Figure VI.3.1a is from a Shuttle Radar Topography Mission (SRTM), NASA/JPL/NIMA picture (taken in February 2000). This project was also funded by a Doctoral grant from the French Ministry of Higher Education and Research and Innovation (Université Paris Cité, Institut de physique du globe de Paris) and ANR V-Care-18-CE03-0010.

ACKNOWLEDGMENTS

This work is part of the PhD thesis of Solaro-Müller (2017).D M would like to thank IPGP for the award of a visiting lectureship which enabled this work to take place. Michel Fialin, Nicolas Rividi and Omar Boudouma are thanked for assistance during EPMA and SEM analyses. We would like to thank the editor V. Acocella, the asociated editor J. Lowenstern and the two reviewers, G. Kilgour and D. Ruth, for their constructive comments.

FUNDING

This work is part of CS (financial support of the French MNRT doctoral Grant). We would like to thank the PPR fund of IPGP. This study was partly founded by the program Tellus-Aleas of INSU (CNRS) and ANR V-CARE-18-CE03-0010.

Supplementary Figure IV.1: Pictures of the outcrops of Grand Fond (GF), Goodwill (GE) and PPR1 to 3 in Dominica. ………………………………………………………………………………...……..166 Supplementary Figure IV.2: Ternary diagram of the unzoned opx compositions of Grand Fond (GF), Goodwill (GE) and PPR1-3. …………………………………………………………………………167 Supplementary Figure IV.3: High resolution Scanning Electron microscope (SEM) images and modelled diffusion profiles for three orthopyroxenes (opx) of our opx set having a reverse (a), a normal zonation (b) and a normal zonation (c) with a bump due to growth near the core in an opx of the PPR1 eruption. ………………………………………………...……………………………………………169 28 Supplementary Figure IV.4: Example of a multiple-zoned orthopyroxene, displaying a third rim measuring less than 10 µm and then not considered as a plateau in the results. ………...…………………………………………………………………………………………..….171 Supplementary Figure IV.5: Individual timescales and their absolute uncertainties based on the propagation of a temperature uncertainty of ± 25 (Solaro et al., 2019) for Grand Fond opx and 9 °C [START_REF]Evidence for an Active, Transcrustal Magma System in the Last 60 ka and Eruptive Degassing Budget (H 2 O, CO 2[END_REF] for the small explosive eruptions and interdiffusion coefficient measurements uncertainties on Grand Fond opx (a), Goodwill (b), PPR1-3 (c-e). ………….. ……………………..172 Supplementary Figure IV.6: En content (%) of the opx hosting the melt inclusions (MI) vs CaO content (wt%) of the MI, represented as a function of the defined magmatic environments (ME) as in Figure 6 (a. Grand Fond (GF); b. Goodwill (GE); c. PPR1; d. PPR2. e. PPR3). ………………………….…..174 Supplementary Figure IV.7: Total Alkali Silica (TAS) diagram for MI in opx. …. ………………….179 Supplementary Figure IV.8: Harker diagrams for MI in opx from this study, plag and cpx and residual glasses. a. K2O vs SiO2; b. Al2O3 vs SiO2; c. CaO vs SiO2 and d. FeO vs SiO2…………..…………183 Supplementary Figure IV.9: Compositions of the unzoned opx compared the cores of the zoned opx in En content (%) vs Al2O3 (wt%) from EPMA analyses (a. Grand Fond,b. Goodwill,c. PPR1,d. PPR2, e. PPR3). ……………………………………………………………………………………………..184 Supplementary Figure IV.10: a. Rank ordering of the interdiffusion timescales modelled in the opx of Grand Fond, Roseau and the small eruptions, as in Figure IV.7.f. ………..…………...…………….186 Supplementary Figure IV.11: Sieve textures in plag crystals of the five eruptions (highlighted by the red arrows; a. Grand Fond, b. Goodwill, c. PPR1, d. PPR2, e. PPR3, red arrows). For plag compositions: see other studies [START_REF]Evidence for an Active, Transcrustal Magma System in the Last 60 ka and Eruptive Degassing Budget (H 2 O, CO 2[END_REF]Solaro et al., 2019). ………………………………………188 Supplementary Figure IV.12: Proportions of the zonations identified in the opx of Grand Fond (a) and the small explosive eruptions (b-e) of the 355 µm fraction (same abbreviations as in Figure 2). ……189 Supplementary Figure IV.13: Proportions of the zonations identified in the opx of the small explosive eruptions (a-d) of the 250 µm fraction. ………………………………………………………...……..191 Supplementary Figure IV.14: Proportions of the zonations identified in the opx of the small explosive eruptions (a-d) of the 125 µm fraction. ……………………………………………………………….192 Supplementary Figure IV.15: Diagram explaining the identification of magmatic environments (ME) from the EPMA compositional profiles after deducing them from the frequency histograms of the En content of all the opx (Figure 3). ……. ……………………………………………………………….193 8: Summary of the crystal assemblages of the blast and Plinian phase, crystal compositions, storage conditions and dynamics of magma storage inferred from crystals zoning (opx or mgt; this study), resorption zones in plag [START_REF] Shcherbakov | Phase equilibria constraints on preeruptive magma storage conditions for the 1956 eruption of Bezymianny Volcano, Kamchatka, Russia[END_REF] or reaction rims in amph [START_REF] Plechov | Opacitization conditions of hornblende in Bezymyannyi volcano andesites (March 30, 1956 eruption)[END_REF][START_REF] Shcherbakov | Phase equilibria constraints on preeruptive magma storage conditions for the 1956 eruption of Bezymianny Volcano, Kamchatka, Russia[END_REF]. ……………………………………………………………..………233 Figure V.9: Magma dynamics inferred for the 1956 climactic phase of Bezymianny, from the combined point of view of opx and mgt of this study and plag and amph from cryptodome clasts and pyroclastic deposits from the Plinian phase (see Figure V.8 Supplementary Figure V.7: (a-d) Harker diagrams for MI in opx of this study and residual glasses in the same samples (a; K2O vs SiO2,b. Al2O3 vs SiO2,c. CaO vs SiO2 and d. FeO vs SiO2) and (e-f) En content (%) of the hosting opx vs CaO (wt%) or Al2O3 (wt%). ……………………………….. …….246 Supplementary Figure V.8: BSE images corresponding to Figure V.3. (a) Unzoned opx; b-c) Singlezoned opx: normal-zoned (b); and reverse-zoned (c); (d-i) Multiple-zoned opx, either normal + reverse (d) or reverse + normal (e) or with three rims: normal + reverse + normal (f) or reverse + normal + reverse (g) or with 4 rims (h) or with more than 5 rims (i). ………………………………………..…247 Supplementary Figure V.9: Proportions of the zonations identified in the opx of the blast in each fraction (500-710, 315-500, 250-355 and 125-250 µm) (a) and in the Plinian phase of the 1956 eruption of Bezymianny in the fractions 315-500, 250-355 and 125-250 µm (b). ………………………………248 Supplementary Figure V.10: Compositions of the unzoned opx compared to the zoned opx (reverse and normal rims and their cores) in the blast deposits and Plinian phase. …………..……………………249 Supplementary Figure V.11: En contents (%) of the zoned and unzoned opx of the blast and Plinian phase. ………………………………………………………….………………………………..……250 Supplementary Figure V.12: Histograms of Fe-Mg interdiffusion timescales modelled in opx of the blast deposits and Plinian phase at 915 ± 17 °C for the blast (a) and 991 ± 44 °C for the Plinian phase (b-c), using the temperature found with the thermometer of Sauerzapf et al. (2008), for core-rim of single-zoned (SZ) opx and inner-outer rim boundaries of multiple-zoned (MZ) opx. d-e) Individual timescales and their absolute uncertainties based on the propagation of a temperature uncertainty of ± 17 °C for the blast or 44 °C for the Plinian phase and uncertainties associated to SEM images (calculated using a Monte Carlo simulation). …………………………………………………………………….251 Supplementary Figure V.13: Individual timescales presented in Figures V. 6-7, modelled in opx (a-b) and mgt (c-d) and their absolute uncertainties based on the propagation of a temperature uncertainty of ± 17 °C for the blast or 44 °C for the Plinian phase and uncertainties associated to SEM images for opx. …………………………………………………………………………………………………..……252 Supplementary Figure V.14: Approximation of the magma storage depths of the Plinian phase of 1956 eruption from CO2 and H2O content of MI hosted in opx, amph, plag or mgt of these samples (MI data from d [START_REF] Ostorero | Time -window into the transcrustal plumbing system dynamics of Dominica (Lesser Antilles)[END_REF]), with isobars according to the rhyolite solubility model of [START_REF] Newman | VolatileCalc: a silicate melt-H2O-CO2 solution model written in Visual Basic for excel[END_REF] and their "VolatileCalc" spreadsheet. ………………………………...…………252 Supplementary Figure V.15: BSE images showing the textures of the rims of amphiboles (in the blast and Plinian phase), with coarse-grained decomposition within a crystal (a-c) or with granular or symplectitic type of rims (d-g). ………………………………………………………………………253

Supplementary Figure V.16: Comparison of the effect of changes in the temperature (T) (a-b), or considering the oxygen fugacity (fO2) or not (c-d) or changes in the diffusion coefficient (D) expression used (e-f) in the timescales modelled, for the blast or Plinian phase. …………………………………254 Supplementary Figure V.17: Distribution of timescales of opx for each dataset according to the changes of some parameters (T, fO2 or not, D), for the blast and Plinian phase (Supplementary Figure V.16) (using the code provided in [START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF]). …………………………………………………255 Supplementary Figure V.18: Distribution of timescales of mgt using the temperature estimated with Sauerzapf et al. (2008) with the diffusion coefficient of [START_REF] Aragon | Cation diffusion in titanomagnetites[END_REF] (using the code provided in [START_REF] Metcalfe | Magmatic Processes at La Soufrière de Guadeloupe: Insights From Crystal Studies and Diffusion Timescales for Eruption Onset[END_REF]). …………………………………………………………………………….256 Figure VI.1.1: From macroscale to microscale: Kizimen location in Kamchatka, 2010-2013 eruption samples and opx zonations. …………………………………………………………………………..286 Figure VI.1.2: Enstatite (En) content histograms of unzoned (a) and single-zoned opx of the 2010-2013 eruption (b) (unzoned opx compositions shown in the dashed frame). ……………………………….287 Figure VI.1.3: Linking monitoring data with timescales of magmatic processes. ……………………288 Figure VI.1.4: Pre-eruptive dynamics of the magmas in the timeframe given by the timescales modelled in zoned opx: crystal clocks mainly starting 1.5 years before the eruption. ……………………........289 Figure VI.1.5: Database showing the comparison between diffusion timescales for studies correlating diffusion timescales and monitoring signals before or during an eruption. ………………………….290 Supplementary Figure VI.1.1: Kizimen volcano in Kamchatka and its geochemical classification. …296 Supplementary Figure VI.1.2: Other textures identified in the opx and proportions of the different zonations in the opx populations in the dacite and andesite (in addition to Figure 1g-j). ……………297 Supplementary Figure VI.1.3: Proportions of the zonations identified in the opx of the dacite in the different fractions (a. 500-710 µm, c. 355-500, e. 250-355 and g. 125-250 µm) and in the opx of the andesite (b. 500-710 µm, d. 355-500, f. 250-355 and h. 125-250 µm). ……………………………..298 Supplementary Figure VI.1.4: Orthopyroxenes En contents (%). …………………………………..299 Supplementary Figure VI.1.5: Compositions of the unzoned opx compared to the zoned opx (reverse and normal rims and their cores) in the dacite and andesite. ………………………………………..300 Supplementary Figure VI.1.6: Histograms of Fe-Mg interdiffusion timescales modelled in opx at 832 ± 37 °C for the dacite (a-b) and 850 ± 57 °C for the andesite (c-d) for core-rim of single-zoned (SZ) opx and inner-outer rim boundaries of multiple-zoned (MZ) opx; (e-f) individual timescales and their absolute uncertainties based on the propagation of a temperature uncertainty of ± 37 °C for the dacite or 57 °C for the andesite and uncertainties associated to SEM images. ……………………………..301 Supplementary Figure VI.1.7: Scanning electron microscope images and En content (blue) (%), Al2O3 (grey) and CaO content (green) (wt%) of zoned opx suitable or not for the modelling of the timescales. ………………………………………………………………………………………………………..302 Supplementary Figure VI.1.8: Examples of measured and modeled opx compositional profiles for three different zoned crystals. ……………………………………………………………………………..303 Supplementary Figure VI.1.9: Rhodes diagrams showing (a-b) opx Mg# for the melt inclusions (MI) hosted in the cores of opx (unzoned or zoned) compared to Mg# of the opx and (c-d) Mg# of the mean residual glasses (RG; andesitic or dacitic) compared to the Mg# of the rims of the zoned opx that should be in equilibrium with the matrix glasses (Supplementary Data VI. 1.1). ……………………………304 Supplementary Figure Harker diagrams for MI in opx of this study (and one plag) and residual glasses in the same samples (a; K2O vs SiO2,b. Al2O3 vs SiO2,c. CaO vs SiO2 and d. FeO vs SiO2) and (e-f) En content (%) of the hosting opx vs CaO (wt%) or Al2O3 (wt%). ………………….305 Supplementary Figure VI.1.11: Cumulative number of the timescales estimated on the major compositional changes evidenced for the andesite and dacite (reverse zoning for the andesite and normal zoning for the dacite; Supplementary Data VI. 1.4) and cumulative number of earthquakes. ……….306 Supplementary Figure images of the homogenized starting materials of the experiment after the melting process in the furnace; a) Synthetic basalt and b) Dacite. …………370 Supplementary Figure VI.3.2: Crystal fractions and viscosities. …………………………………….370 Supplementary Figure VI.3.3: 2D Microlite number density versus the crystal fraction crystal microlite content φ (a) and (b-d), orientations of the crystals for the experiment, or for the natural banded lavas, either in the andesitic band or the dacitic band. ……………………………………………………..371 Supplementary Figure VI.3.4: Evolution of the sizes of the crystals depending on their location in the experimental sample. ………………………………………………………………………………...372 Supplementary Figure ) Compositions of the residual glasses of the crystal enclaves of the experiment for major elements. ……………………………………………………………………...373 Supplementary Figure VI.3.6: Compositions of the experimental plagioclase crystals and clinopyroxenes analyses. …………………………………………………………………………….374 Supplementary Figure VI.3.7: Chemical maps of the experimental plagioclase crystals: a and e) in grey intensities, b and f) Ca mapping, c and g) Na mapping and d and h) Ca, Na and K maps in Red-Green-Blue (RGB). ………………………………………………………………………………………….375 Appendix Figure A.2: From crystal to greyscale profile. …………………………………………….463 Appendix Figure A.3

LIST OF TABLES AND SUPPLEMENTARY TABLES

Supplementary Table IV.1: Excel file with major element analyses (wt%), En content and Mg number (Mg#) in zoned opx of Grand Fond, Goodwill, PPR1, 2 and 3 (in the three fractions) and all associated profiles in En content vs distance from the rim (µm), as well as major element analyses (wt%) in unzoned opx of Grand Fond, Goodwill, PPR1-3, En content and Mg number (Mg#). ………………195 Supplementary Table IV.2: En content peaks of the unzoned opx and major En change depending on the two populations of opx for the five eruptions, according to Figure IV.3 and 4. SZ: single-zoned opx, MZ: multiple-zoned opx. …………………………………………………………………………….196 Supplementary Table IV.3: Excel file with timescales estimates for Grand Fond (for a temperature of 850°C (Solaro-Müller, 2017;Solaro et al., 2019)), Goodwill, PPR1-3 eruptions (for a temperature of 890°C [START_REF]Evidence for an Active, Transcrustal Magma System in the Last 60 ka and Eruptive Degassing Budget (H 2 O, CO 2[END_REF]) with references to the different pathways and ME changes recorded (name of the samples: DOM43a1 for Grand Fond, DOM48c2 for Goodwill, DOM51b2 for PPR1, DOM49d2 for PPR2 and DOM49i2 for PPR3). …………………………………………………….197 Supplementary Table IV.4: Percentages of magma volumes in the two ME for each eruption, estimated by the proportions of unzoned and zoned opx (on the one hand, from the proportions of Figure 2 and on the other hand, from the histograms in Figure 3 and4: calculated using the number of single-zoned (SZ) and multiple-zoned (MZ) opx of the two populations). ………………………………………..198 Supplementary Table IV.5: Major element analyses (wt%) in MI and hosts (opx) of Goodwill (GE), PPR1, 2 and 3 in the 355 and 250 µm fractions (name of the samples: DOM48c2 for Goodwill, DOM51b2 for PPR1, DOM49d2 for PPR2 and DOM49i2 for PPR3). ……………………………..199 Table V.1: Description of the 1956 climactic phase samples of Bezymianny studied here. …………211 Table V.2: Temperatures estimated using magnetite/ilmenite pairs [START_REF] Ghiorso | Thermodynamics of Rhombohedral Oxide Solid Solutions and a Revision of the FE-TI Two-Oxide Geothermometer and Oxygen-Barometer[END_REF]Sauerzapf et al., 2008;[START_REF] Yavuz | WinMIgob: A Windows program for magnetite-ilmenite geothermometer and oxygen barometer[END_REF]. …………………………………………………………………………...219 Appendix Table V.1: Summary of the timescales results. Median of the timescales in opx, expected timescale and its standard error associated for each dataset are shown, according to the changes of some parameters (T, fO2 or not, D), for the blast and Plinian phase, as shown in Supplementary Figure V.13, S15-16. ………………………………………………………………………………………………238 Supplementary Table VI.1.1: Description of the 2010-2013 samples from Kizimen eruption studied here (the localization, coordinates and nature of the samples are specified). …………………………306

Table VI.2.1: Description of the four types of plagioclase and their proportions in andesites and dacites, with the same colour code as in Figure VI.2.3. ………………………………………………………325 Table VI.2.2: Timescales modelled in plag of the andesite. Name of the crystals, zonation types (NZ: normal-zoned, RZ: reverse-zoned and errors associated to the plag are specified (corresponding to the modelling with the uncertainties on the temperature considered). …………………………………..332 Table VI.3.1: Major elements composition of the basaltic enclave of Kizimen (Churikova et al., 2013) that was reproduced through the synthetic basalt and the starting material compositions of the synthetic basalt and dacite, with the positive and negative errors. …………………………………………….347 Table VI.3.2: Experimental conditions adopted during magma mixing experiments. ……………….350 Table VI.3.3: CSD results. …………………………………………………………………………...354 Table VI.4.1: Summary of the experimental conditions and the temperatures of the first melting process (first at 1500 °C), and then at the temperature where a crystal fraction of 13 vol% of feldspar may be generated at hypothesized NNO + II oxidizing conditions (simulated with MELTS) [START_REF] Gualda | Rhyolite-MELTS: a Modified Calibration of MELTS Optimized for Silica-rich, Fluid-bearing Magmatic Systems[END_REF]. ………………………………………………………………………………………………...378 Table VI.4.2: Whole rock starting material compositions (dacite and basaltic andesite Kizi-96-24 from Churikova et a. ( 2013)) and the natural basaltic andesite before the first melting process , analyses carried out at CRPG, as in Chapter VI. 3

LIST OF SUPPLEMENTARY DATA

Supplementary Data V.1: Whole rock major elements (see the TAS diagram, Figure V.2) and major elements compositions (wt%) in melt inclusions, host minerals in the blast deposits and Plinian phase and residual glasses in the blast deposits. …………………………………………………………….257 Supplementary Data V.2: Excel file with Fe-Ti magnetite/ilmenite (mgt/ilm) pairs major element analyses (wt%) and associated temperatures estimated using the mgt/ilm thermometer [START_REF] Ghiorso | Thermodynamics of Rhombohedral Oxide Solid Solutions and a Revision of the FE-TI Two-Oxide Geothermometer and Oxygen-Barometer[END_REF]Sauerzapf et al., 2008;[START_REF] Yavuz | WinMIgob: A Windows program for magnetite-ilmenite geothermometer and oxygen barometer[END_REF]. …………………………………………………257 Supplementary Data V.3: Excel file with whole major element analyses (wt%) of opx, En content and Mg number (Mg#) in zoned opx of the blast and Plinian phase of the 1956 eruption of Bezymianny (in the four fractions) and all associated profiles in En content vs distance from the rim (μm), as well as major element analyses (wt%) in unzoned opx of the blast and Plinian phase, En content and Mg number (Mg#). ………………………………………………………………………………………………..257 Supplementary Data V.4: Excel file with timescales estimates modelled in the zoned opx of the blast and Plinian phase (using average temperatures of Supplementary Data V.2, changing the oxygen fugacity or the diffusion coefficient parametrization), with references to the different pathways and En changes recorded. ……………………………………………………………………………………257 Supplementary Data V.5: Excel file with Fe-Ti magnetite (mgt) whole major element analyses (wt%) profiles and the timescales modelled in the mgt of the blast and Plinian phase. Timescales have been modelled with the temperatures estimated by Sauerzapf et al. (2008)

Supplementary Data VI. 1.1: Whole rock major elements (see the TAS diagram, Supplementary Figure 1) and major elements compositions (wt%) in melt inclusions, host minerals in the dacite and andesite and residual glasses in the dacite, andesite and banded lavas. ………………………………………306 Supplementary Data VI. 1.2: Excel file with major element analyses (wt%), En content and Mg number (Mg#) in zoned opx of the dacite and andesite of the 2010-2013 eruption of Kizimen (in the four fractions) and all associated profiles in En content vs distance from the rim (µm), as well as major element analyses (wt%) in unzoned opx of the dacite and andesite, En content and Mg number (Mg#).

………………………………………………………………………………………………………..307

Supplementary Data VI. 1.3: Excel file with Fe-Ti magnetite/ilmenite (mgt/ilm) pairs major element analyses (wt%), En content of the dacite and andesite of the 2010-2013 eruption of Kizimen (in the 250 µm fraction) and associated temperatures estimated using the mgt/ilm thermometer4. …………….307 Supplementary Data VI. 1.4: Excel file with timescales estimates modelled in the zoned opx for the dacite (for a temperature of 832 °C) and andesite (for a temperature of 850 °C) with references to the different pathways and En changes recorded (name of the samples: Kizi 5, 8 and 9 for the andesites and Kizi 6 for the dacite). …………………………………………………………………………………307 Supplementary Data VI. 1.5: Summary of the database of studies linking petrological signals (diffusion timescales) with monitoring signals, either pre-or syn-eruptive signals, shown on Figure VI.1.5. ..307 Supplementary Data VI.2.1: Excel file with major element analyses (wt%) of the plag and mgt and Ti (wt%) of the profiles done in the zoned mgt of the andesite and dacite. The summary of all the profiles is also shown. ………………………………………………………………………………………...339 Supplementary Data VI. 2.2:

Competing interests

The authors declare no competing interests.

165

Time-window into the transcrustal plumbing system dynamics of Dominica (Lesser Antilles)

Lea Ostorero

Supplementary material

Grand Fond opx (a), Goodwill (b), PPR1-3 (c-e). These uncertainties are due to the diffusivity calculations (D0, the interdiffusion coefficient, the activation energy Ea), temperature and resolution of grayscale values of the BSE images 5,6 and are calculated thanks to a Monte Carlo simulation. These uncertainties are asymmetric, with a larger positive error bar and a smaller negative error bar, due to the logarithmic effect of the uncertainties on mainly temperature and diffusivity calculations 7 . The error bars show a high variability, especially for the longest timescales, which places a limit on the statistics of the modelled timescales and their significance.

Supplementary Note: previous results: compositions of natural products and pre-eruptive magma storage conditions

Mineralogy

Samples from the five eruptions (Grand Fond, Goodwill, PPR1, PPR2, PPR3) have been studied in terms of whole-rock compositions, crystal content, melt inclusions (MI) and residual glass compositions in major and trace elements 4,9,10 . Grand Fond and the following small explosive eruptions magmas have whole rock compositions between acid andesites and dacites, with 63 wt% SiO2 and ~4.5-5 wt% Na2O+K2O for Grand Fond 3,9,10 (recalculated with total iron as FeO and on anhydrous basis) and 62-63 wt% SiO2, ~4.5-5 wt% Na2O+K2O for the small explosive eruptions 4 (Supplementary Figure S7). All the residual glasses are rhyolitic wt% Na2O+K2O) 4,9,10 (Supplementary

Figure S7).

Large explosive eruption: Grand Fond Grand Fond pumices contain ~30 vol% of phenocrysts: plagioclase (plag, ~21 vol%), orthopyroxene (opx, ~5 vol%) and smaller proportions of clinopyroxene (cpx) and Fe-Ti oxides (mgt) (<1 vol%) 3,9 . The anorthite content of the plag varies from An46 to An94 with rims of An50-55. Opx are hypersthenes; they span a broad enstatite range, from En49 to En56 with ~80-85% of unzoned crystals and cpx are diopsides with Wo44-45 and En37-38 3 . Fe-Ti oxides (mgt) show compositions with ~75-80 wt% FeOtot and ~10 wt% TiO2 (Mag71-75) 3 .

Small explosive eruptions: Goodwill, PPR1 to 3

Pumices from the small explosive eruptions have a similar crystal assemblage as Grand Fond, in decreasing order of abundance: plag >> opx >> cpx >> mgt 4 . For PPR1 and PPR3, amphiboles are present in low proportions. The anorthite content of plag varies from An42 to An70, the enstatite content for opx is between En46 to En59, and the wollastonite content for cpx varies from Wo38 to Wo44. Mgt compositions are between Mag69 and Mag91 4 .

Major elements composition of melt inclusions and matrix glass

Grand Fond: All MI analyzed in opx and plag exhibit a rhyolitic composition and are located in the calc-alkaline field in the K2O vs SiO2 diagram (with Al2O3 contents between 11 and 15 wt% and low CaO content, between 1 and 2.5 wt%; Supplementary Figure S8) 10 . For opx, their Na2O+ K2O content varies from 4.9 to 6.5 wt% and their SiO2 content from 75.6 to 79 wt%. For plag, they show variations in Na2O + K2O of 5.0-5.6 wt% for a SiO2 range of 76.1 to 78.2 wt%. The residual glasses show a SiO2 content between 76 and 77 wt% and a Na2O + K2O content between 5 and 6.5 wt% 10 .

Small explosive eruptions: All MI trapped in opx, cpx, plag and mgt have a rhyolitic composition (71-79 wt% SiO2, 3.5-7 wt% Na2O+K2O, Supplementary Figure S7) 4 . In the K2O vs SiO2, the MI in all the host minerals are in the medium-K subalkali compositional field Supplementary Figure S10: a. Rank ordering of the interdiffusion timescales modelled in the opx of Grand Fond, Roseau and the small eruptions, as in Figure 7.f. Timescales modelled on Grand Fond's bumps were added for comparison (Supplementary Table S3). b. Individual timescales and their absolute uncertainties based on the propagation of a temperature uncertainty of ± 25 °C4 and interdiffusion coefficient measurements uncertainties on Grand Fond opx. These uncertainties are due to the diffusivity calculations (D0, the interdiffusion coefficient, the activation energy Ea), temperature and resolution of grayscale values of the BSE images 5,6 and are calculated thanks to a Monte Carlo simulation. These uncertainties are asymmetric, with a larger positive error bar and a smaller negative error bar, due to the logarithmic effect of the uncertainties on mainly temperature and diffusivity calculations 7 . The error bars show a high variability, especially for the longest timescales, which places a limit on the statistics of the modelled timescales and their significance.

Tables

Supplementary Table S1:

En53

Supplementary Table S3: Excel file with timescales estimates for Grand Fond (for a temperature of 850°C 3,8 ), Goodwill, PPR1-3 eruptions (for a temperature of 890°C The same En contents are then identified for both the unzoned opx and cores of the zoned opx in the blast and Plinian phase and the same main En variations are identified in the rims of the zoned opx of both deposits. In both deposits, regarding the multiple-zoned opx and their last rims (called "N"), two populations of last rims can be identified: last normal-zoned rims (going from En65-68 (N-1 rim) to En58- V. 4.6.3. Timescales Changing of the critical parameters for timescale modelling, such as temperature T following the geothermometer used or adding or not the effect of oxygen fugacity fO2 or changing the diffusion coefficient expression, have effects on the timescales modelled and are discussed in the Appendix at the end of the paper (Supplementary Data V.4). Here, timescales are modelled using the same method of temperature estimation (915 ± 17 °C for the blast and 991 ± 44 °C for the Plinian phase (Sauerzapf et al., 2008;[START_REF] Yavuz | WinMIgob: A Windows program for magnetite-ilmenite geothermometer and oxygen barometer[END_REF] and using Ganguly and Tazzoli (1994) without the fO2, as the opx are in the appropriate En ranges and that the fO2 effect is not constrained (Methods; Table V.2; Supplementary Figure V.12; Supplementary Data V.3-4).

Blast

Out of 38 transects in zoned crystals, only the profiles with sigmoidal shapes (due to diffusion), were kept, thus, 30 profiles were modelled in the opx of the blast (Figure V.6; Supplementary Figure V.12; Supplementary Data V.3-4). The timescales estimated on all the rims are mainly below 1 year before the eruption, with a majority of timescales below 200 days (~6 months) (Supplementary Figure V.12). The timescales have been modelled on both inner rims and external rims with the same temperature (last rim N, N-1 and N-2…). The inner and N-1 rims give timescales from 3 years to 15 days before the eruption, with a majority of normal-zoned opx that are below 6 months (Figure V.6a-b). The timescales of the last rims N of zoned opx range from 6 months to 4 days before the eruption, with the majority of the timescales that are below one month for reverse-zoned rims, prior to the climactic phase (Figure V.6a-b; Supplementary Data V.4). For timescales below 200 days, the uncertainties are between + 1000/ -3 days (Supplementary Figure V.13; Supplementary Data V.4). For timescales from 1 year to 3 years, the uncertainties are between + 4000/ -190 days (Supplementary Figure V.13; Supplementary Data V.4).

Plinian phase

Out

V.4.7. Mgt

V. 4.7.1. Zonations 1281 mgt were mounted, in the fractions 315-500 and 250-315 µm (637 for the blast and 644 for the Plinian phase). Mgt in both samples are all reversely-zoned, with rare ones being multiple-zoned.

SUPPLEMENTARY MATERIALS FOR CHAPTER V:

Crystals record of magma dynamics before the 1956 flank collapse eruption of Bezymianny

Provisional list of authors: Lea Ostorero 1* , Georges Boudon 1 , Hélène Balcone-Boissard 2 , Thiebaut d'Augustin 2 (rest of the author list to be defined) 1 Université Paris Cité, Institut de physique du globe de Paris, CNRS, F-75005 Paris, France

CHAPTER VI: KIZIMEN PRE-ERUPTIVE TIMESCALES THROUGH CRYSTAL ZONING AND EXPERIMENTAL STUDIES

used to model the timescales are calculated on magnetite/ilmenite couples in the dacite and andesite, this equilibrium between these minerals has been widely used to estimate temperatures in igneous rocks 47 . Mean temperatures of 832 ± 37 °C and 850 ± 57 °C are found for the andesite and dacite respectively 47 (Supplementary Note, Supplementary Figure 6 and Supplementary Data 3-4). These profiles have sigmoidal shapes due to diffusion effects, without bumps or peaks that could indicate a growth component 26 (Supplementary Figure 78and Supplementary Data 2). Al2O3 and CaO gradients are also compared to the En content to verify that the gradients in En are due to diffusion 48 (Supplementary Figure 7; Supplementary Data 2). For both the dacite and andesite, most of the timescales are young (< 2 years) for single-zoned and outer rims of multiple-zoned opx (Supplementary Figure 6, Supplementary Note and Supplementary Data 4). Longer timescales associated to multiple zonations can be related to previous events in the magmatic reservoir, recorded by the opx. On the major En changes recorded by the zoned opx (En62-64 to En64-73 in the reverse zoned opx of the andesite and from En62-64 to En44-54 in the normal-zoned opx of the dacite), 80% of the timescales modelled are in the 1.5-years period prior to the eruption whatever the type of zoning (normal or reverse) (Figure 3b andd; Supplementary Figure 6 and8 and Supplementary Data 4).

Discussion

Magmatic processes modify the storage conditions of magmas before volcanic eruptions (changes in temperature, pressure, oxygen fugacity (fO2) or volatile content) 7,13,15,26,49-51 and Fe-Mg interdiffusion timescales modelling of the zonations in the opx help us to retrieve the timescales of these processes by studying the start of the "clocks" of these crystals (Figure 3b andd; Supplementary Figure 6 and Supplementary Data 4).

70% of the opx are unzoned (Supplementary Figures 23). Different hypotheses can be evoked to explain this high proportion of unzoned opx: i) the magmatic changes did not affect the entire magmatic reservoir; ii) the unzoned crystals were older and their zoning patterns were homogenized by diffusion with time; iii) they are formed after the last magmatic perturbation discussed below. The size of the unzoned crystals and their proportions compared to zoned opx are comparable in the different fractions studied (from 125 to 710 µm; Supplementary Figures 23). Their compositions are also in the same range as the cores of the zoned opx, in En62-64 and in the same domains in the CaO, FeO, MgO and MnO versus SiO2 diagrams, compared to the rims compositions of the zoned opx (Figure 2; Supplementary Figures 45). It was then more likely that they formed before the last magmatic event prior to the eruption and that the magmatic changes did not affect the entire magmatic reservoir (Figure 3a).

As all the unzoned and zoned opx have the same core compositions (Supplementary Figure 45and Supplementary Data 2), the opx from the dacite and the andesite are likely to be originated from the dacitic magma present in the reservoir before the 1963 basaltic injection. The andesitic part of the reservoir was then probably formed after the injection of an aphyric basaltic magma in 1963 and mixed with the dacite in the lower part of the reservoir, as suggested in a previous study of the eruption 41 and incorporated some of the crystals of the dacite (including opx) (Figure 4a). This is also confirmed by a previous study on plagioclase crystals in banded rocks of the same eruption 41 .

The majority of the zoned opx studied indicate a short timescale (< 1.5 years) and no zoned opx reflects a longer timescale in the order of 50 years (that could correspond to a possible mafic injection in 1963) (Supplementary Figure 6).

The uncertainties associated to the timescales modelled are mainly due to the temperature uncertainty and uncertainties due to diffusion coefficient determination 6,26,52 (see Methods; Figure 2

Competing Interest Statement

All the authors declare no competing interests.

SUPPLEMENTARY MATERIALS FOR CHAPTER VI.1:

Insights into volcanic unrest by correlating petrological and seismic observations at Kizimen volcano (Kamchatka, Russia)

Lea Ostorero 1). The compositions of the samples are in the same domain as in a previous study 41 (Supplementary Figure 1). Samples are porphyritic: ~36% of crystals for the andesite and ~29% of crystals for the dacite (the crystallinities of the samples have been estimated after a vesiclefree estimation by point counting on thin sections of the andesite and dacite samples thanks to Mesurim (Version 3.4; http://acces.ens-lyon.fr/acces/logiciels/applications/mesurim)).

Residual glasses and melt inclusions (MI).

MI trapped in opx in the andesite have a rhyolitic composition (77-80 wt% SiO2, 3.3-6 wt% Na2O+K2O, Supplementary Figure 1; Supplementary Data 1). The only plag-hosted melt inclusion in the dacite also has a rhyolitic composition (Supplementary Figure 1; Supplementary Data 1). For opx, we used the Rhodes' diagram to test for equilibrium for orthopyroxene, using the following equation:

Where 𝑋 𝑆𝑖 𝑙𝑖𝑞 is the cation fraction of SiO2 in the equilibrium liquid. The mean KD(Fe-Mg) opx-liq for the five inclusions is KD(Fe-Mg) opx-liq = 0.19±0.05, which implies < 1% of recrystallization. For the one MI in plag, the measured KD(An-Ab) plag-liq = 0.12, which implies ~7% of recrystallization. The Rhodes diagram in Supplementary Figure 9 shows the Mg# of the liquid (measured in the melt inclusion) compared to the Mg# of the orthopyroxene 57 .

The residual glasses also plot in the domain of the melt inclusion compositions, with andesitic and dacitic ones in the same domain [START_REF] Ruprecht | Feeding andesitic eruptions with a high-speed connection from the mantle[END_REF][START_REF] Ruth | Crystal and melt inclusion timescales reveal the evolution of magma migration before eruption[END_REF][START_REF] Albert | Years to weeks of seismic unrest and magmatic intrusions precede monogenetic eruptions[END_REF][START_REF] Passarelli | The correlation between run-up and repose times of volcanic eruptions[END_REF][START_REF] Hartley | Tracking timescales of short-term precursors to large basaltic fissure eruptions through Fe-Mg diffusion in olivine[END_REF] wt% Na2O+K2O, Supplementary Figure 1) and residual glasses in the banded lavas with the dark glass that is slightly less evolved than the light glass (75-79 wt% SiO2, 5.6-7 wt% Na2O+K2O, with one point at 71 wt% SiO2, 6 wt% Na2O+K2O and 77.5-79 wt% SiO2, 5.9-6.7 wt% Na2O+K2O, respectively; Supplementary Figure 1; Supplementary Data 1), as observed in a previous study on residual glasses of banded lavas from the 2010 eruption of Kizimen 41 . In the K2O, Al2O3, CaO and FeO versus SiO2 diagrams, the compositions of the MI and residual glasses are in the same domain as the banded lavas from another study 41 and melt inclusions from the fourth eruptive cycle of Kizimen 120 (Supplementary Figure 10). No significant difference between MI hosted in zoned opx or unzoned opx is present, with the En content of the hosted opx between En62-64 and CaO and Al2O3 contents between 0.4-2.4 and 10-13 wt%, respectively (Supplementary Figure 10).

Orthopyroxene zonations: types and proportions. 1841 orthopyroxene (opx) crystals were studied for the 2010-2013 eruption (Supplementary Figure 23). We analyzed them using a scanning electron microscope and electron probe microanalyzer (EPMA). Zoned crystals were classified according to the type of zonation observed: either a single one (normal or reverse zonation) or a multiple one (Supplementary Figure 23). In both dacite and andesite, the proportions of unzoned crystals are higher than zoned ones (66% and 67% of unzoned crystals respectively; Supplementary Figure 2). The proportions of unzoned and zoned opx are very similar whatever the considered fraction (from 59 to 70% of unzoned opx for the dacite and from 53 to 75% for the andesite; Supplementary Figure 3). Considering all the four fractions (500-710, 355-500, 250-355 and 125-250 μm fraction), 31-32% of them are zoned, with a majority of normal-zoned opx in the dacite while reverse-zoned opx in the andesite are more represented (Supplementary Figure 2). Multiple-zoned opx represent a small proportion of zoned crystals (3-6%) and are mainly reverse-zoned then normal-zoned, from core to rim (Supplementary Figure 2). Some opx also show reactional borders or sieved textures and represent 2 % of the opx population (Supplementary Figure 2).

Orthopyroxene zonations: compositions. Profiles in zoned opx were mainly acquired in single-zoned opx (normal-zoned opx for the dacite and reverse-zoned opx for the andesite) and some multiple-zoned opx with the EPMA (37 for the dacite and 75 for the andesite) (Figure 2, Supplementary Data 2 and Supplementary Figure 4). Unzoned opx compositions were also acquired (41 for the dacite and 67 for the andesite) (Figure 2, Supplementary Data 2 and Supplementary Figure 4). The range of compositions displays no systematic differences for the cores of the unzoned and zoned opx of both samples (En62-64; Enstatite content (𝐄𝐧 = 𝐌𝐠 𝐌𝐠+𝐅𝐞

)). The singly normal-zoned rims of the dacite show lower Enstatite content (En44-54) compared to singly reverse-zoned rims with higher En content of the andesite opx (En64-73) (Figure 2, Supplementary Data 2). For multiple-zoned ones, they mainly first show a reverse rim (En64-70 for the andesite vs En64-66) and then a normal rim (En58-70 for the andesite vs En61-64 for the dacite) (Supplementary Data 2; Supplementary Figure 4). Some rims in the multiple-zoned opx of the dacite are normal-zoned and then reverse-zoned (En42-46 to En48-52) (Supplementary Data 2; Supplementary Figure 4). Microlites (between 10-100 µm long and 5-10 µm wide) in the andesitic samples are between En63-72 (Supplementary Data 2). Cores of opx with sieved textures in the dacite have the same compositions as the other opx (En62) while the ones in the andesite have compositions from En63-65 (Supplementary Data 2). In the CaO, FeO, MgO and MnO versus SiO2 diagrams, the cores of zoned opx and unzoned opx are in the same domain (51)(52)(53)(54)(21)(22)(23) wt% MgO and 1.1-1.34 wt% MnO; Supplementary Figure 5). Unzoned compositions in CaO and Al2O3 (wt%) are also in the same domains as the zoned cores (Supplementary Figure 5).

Timescales. Timescales have been modelled on the most representative opx zonations based on

Supplementary Figure 2. For the dacites, timescales are around 1-6 years for the normal-zoned crystals, with 80% of them between 1-2 years (Supplementary Figure 6). For the multiple-zoned opx, the bulk of the timescales are also younger than 2 years (for the inner rims (reverse or normal-zoned): (from En62-64 to En44-54, and for the outer rims of other crystals: timescales between 1-5 years (in the same En range En44-54) and sometimes longer timescales (2 timescales: ~25 years, from En62-64 to En64-73 or in En62-64) (Supplementary Figure 6; Supplementary Data 4). For the andesites, the timescales are distributed between 1 to 4 years for the single reversely-zoned crystals, with 90% of them between 1-2 years before the eruption (Supplementary Figure 6). Normalzoned crystals give two timescales between 1 to 11 years (Supplementary Figure 6). For the multiplezoned crystals, the bulk of the analyses also represent the final 2 years pre-eruption, with little uncertainties (3 inner rims between 1-8 years and one inner reverse-zoned rim gives a timescale of 62 years; for the outer ones, 4 reverse-zoned ones are ~1 year and two normal ones on different multiplezoned crystals than the ones on which core to inner rims were modelled are between 2-23 years; Supplementary Figure 6; Supplementary Data 4). The timescales modelled on zonations in the same En range (En62-64, as the cores of the opx) are between 0-22 years (Supplementary Figure 6; Supplementary Data 4).

Magma temperatures estimation.

Equilibrium temperatures were calculated from the compositions of touching pairs of titanomagnetite and ilmenite (mgt/ilm) following a numerical model 47 (Supplementary Data 3). A mean temperature of 832 ± 37 °C and 850 ± 57 °C were calculated for the dacite (on 3 pairs) and andesite (on 21 pairs) respectively (where the number that follows the ± sign is a standard deviation; Supplementary Data 3). A temperature of 874 °C has also been estimated thanks to the plag/liquid thermometer (equation 24a) 57 , knowing an estimation of the pressure 41 (~123 MPa), as well as host and melt inclusion compositions (see Supplementary Data 1). The estimated water content is of 2.4 wt% H2O, using the hygrometer from a study 58 . From previous results, banded lavas yield H2O values between 5-5.7 wt% H2O in the dacitic bands and 4.9-5.6 wt% H2O in the andesitic bands 41 from amphibole thermometers 64,65 and composition is given) 120 , for the 7550 14 48,101,113,114 Supplementary Table 1. Description of the 2010-2013 samples from Kizimen eruption studied here (the localization, coordinates and nature of the samples are specified). Supplementary Data 1. Whole rock major elements (see the TAS diagram, Supplementary Figure 1) and major elements compositions (wt%) in melt inclusions, host minerals in the dacite and andesite and residual glasses in the dacite, andesite and banded lavas. The sampling locations are specified in Supplementary Table

Abstract

Textures and compositional zoning in minerals from volcanic rocks are witnesses of magmatic processes that occurred during their growth histories. Here, we examined three types of crystals in the crystal assemblage of the magmas emitted during the first historical magmatic eruption of Kizimen in 2010 (Kamchatka). These three crystals (plagioclase, orthopyroxenes and magnetites) are sensitive to storage and emission conditions variations in the magma plumbing system, such as temperature changes and can give us information on the pressure, temperature, and melt composition. By combining a Crystal System Analysis approach and timescales modelling on the zoning patterns of these three crystal types, we reconstruct the pre-eruptive magma dynamics, involving processes on different timescales. A first mixing event following a mafic injection in the dacitic reservoir at depth was identified and occurred more than 5 years before the eruption. This event is responsible for the creation of a temperature-zoned reservoir, composed of a remnant dacitic magma at the top and a newly created andesitic magma at the base. A second mixing process, at the interface between the dacitic and the andesitic magmas, mainly occurred 1.5 years prior eruption and is correlated to the onset of a seismic crisis. Finally, as the closure temperature of magnetites is lower than the one of orthopyroxene or plagioclase, magnetites may have recorded a heating process in the magmatic reservoir or during magma ascent/stay in the lava dome. The magma dynamics unraveled here thanks to the combination of several minerals and timescales modelling show the potential of this type of study for deciphering the magma plumbing history before an eruption.

Keywords: plagioclases, orthopyroxenes, magnetites, timescales, magma dynamics

VI.2.1. Introduction

Recently, to explain the petrological and geochemical diversity observed within volcanoes, the paradigm describing the magma plumbing system has evolved towards a multi-depth magma accumulation through the crust referred as a "transcrustal magma system" (Bachmann and Huber, 2016;Cashman et al., 2017;[START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF]. Former permanent magma chambers are now viewed as ephemeral magma accumulation within a mush, consisting in silicate melt evolving through a crystalline framework (Bachmann and Bergantz, 2006;Barker et al., 2016;Caricchi and Blundy, 2015;[START_REF] Hildreth | Compositional Zoning of the Bishop Tuff[END_REF]Marsh, 1989[START_REF] Marsh | On the crystallinity, probability of occurrence, and rheology of lava and magma[END_REF][START_REF] Petrelli | Rates and Timescales of Magma Transfer, Storage, Emplacement, and Eruption[END_REF]. Magma ponding zones are open systems created by magma supply from depth operating with crustal deformation and rapid segregation of melt and associated fluids in mushy zones [START_REF] Sparks | Formation and dynamics of magma reservoirs[END_REF].

On active and dormant volcanoes, what can be of interest to better manage a potential volcanic crisis is to know the timescales separating the beginning of the unrest detected on a volcano from the onset of the eruption. To do so, eruption records (from the petrology of the deposits, geophysical and geochemical monitoring if they exist) may be correlated as successfully recently demonstrated [START_REF] Liu | Petrologic monitoring at Volcán de Fuego, Guatemala[END_REF]. This opens the way to petrological monitoring, which is particularly interesting on active volcanoes with no monitored eruption (Re et al., 2021). The development of this approach is possible thanks to the increasing data on intracrystalline elemental diffusivity in a large set of crystals (Chakraborty, 2008;Costa, 2021;Costa et al., 2020;Costa and Morgan, 2010;Petrone and Mangler, 2021;[START_REF] Zhang | Diffusion in Minerals and Melts: Introduction[END_REF] and in parallel, to the increasing performances of point analyses allowing to unravel the diversity of crystal textures. Indeed, minerals are formed from magmatic liquids The morphology of Kizimen is a stratocolcano made of lava domes and thick lava flows, surrounded by volcanoclastic material, with its base being covered by pyroclastic density current deposits intersected by a network of valleys (Melekestsev et al., 1995). (Melekestsev et al., 1995).

VI.2.2.2. Kizimen eruptive history

The volcano was formed during the Upper Pleistocene-Lower Holocene period, with the reactivation of pre-existing late Pleistocene magma chambers following the formation of the Shchapinsky graben and the Tumrok ridge horst (Melekestsev et al., 1995). Four cycles of activity have been identified at Kizimen, dated from 12-11 ka to the present day, prior to the recent eruption of 2010, through a tephrochronological study (Melekestsev et al., 1995). They correspond to the period between 12,000-11,000 and 8,400 years ago, 8,400 and 6,400, 6,400 and 3,000 and from 3,000 years ago to the present (Melekestsev et al., 1995) . Each cycle is characterized by periods of high activity beginning with large explosions and ending with the formation of an extrusive lava dome and the emission of lava flows, followed by a phase of decline and finally a phase of rest (Melekestsev et al., 1995). These cycles mostly differ in the volumes and compositions of erupted rocks (variation from dacite to andesite-basalt due to mixing of dacitic and basaltic magmas), in eruptive parameters and in magmatic productivity. Through its eruptive history, Kizimen was a source of heterogeneous, poorly mixed magmas with compositions ranging from dacites to basaltic andesites (Auer et al., 2018;Melekestsev et al., 1995;[START_REF] Piip | Kizimen volcano[END_REF].

In the last eruptive cycle, the volcano produced an eruption in 1927-1928, probably of hydrothermal origin; no deposit was found but a strong earthquake which occurred on 21 February 1928 in the area of Kizimen was recorded (Auer et al., 2018;Churikova et al., 2007;Melekestsev et al., 1995). A fumarole has been continuously present since 1825 at the north-western slope of the volcano at 1950 m a.s.l., 400 m below the volcano summit as witnessed by local residents (Auer et al., 2018;[START_REF] Kirsanova | Fumarolic activity of Shiveluch and Kizimen volcanoes in 1979-1980[END_REF][START_REF] Piip | Kizimen volcano[END_REF]. In 1963, the volcano was the site of intense seismic activity (earthquakes recorded from the ground surface up to 35 km deep) [START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF]. In 2010-2013, the volcano produced its first historical magmatic eruption of magma with explosive episodes generating pyroclastic density currents followed by the extrusion of a lava dome that gradually formed a very thick lava flow (Auer et al., 2018).

Numerous block and ash pyroclastic density currents were produced resulting from the gravity collapse of the lava dome and lava flow front (Auer et al., 2018). There were no human settlements near Kizimen and the eruption did not cause any casualties.

VI.2.2.3. The 2010 eruption: magmatic activity

After 92 years of repose, an explosive-effusive eruption occurred in December 2010, after 18 months of slow-going seismic precursory process (Firstov and Shakirova, 2014;[START_REF] Senyukov | Seismicity of Kizimen volcano[END_REF][START_REF] Shakirova | Features of the Kizimen Volcano area seismicity prior to and during the 2010-2013 eruption[END_REF].

VI.2.4.3. Mineral compositions

Orthopyroxenes and plagioclase compositions are described together, from core to rim. VI.2.4.3.1. Cores Opx: The compositions of the cores of zoned opx and unzoned ones have a restricted compositional range of En62-64 (Figure VI.2.4a) (Ostorero et al., accepted).

Plag: Whatever the emitted magma, dacitic or andesitic, the composition of the plag cores describe a wide range of compositions, An30-90

VI.2.7. Supplementary Data (available upon request)

Supplementary Data VI.

VI.2.8. Acknowledgements

We would like to thank S. Hidalgo for her help with sample preparation, M. Fialin and N. Rividi for assistance during EPMA analyses and O. Boudouma and S. Borensztajn for SEM imaging. We also thank G. Fabbro for his precious help on the diffusion code modelling of plag. This project was funded by the Institut de physique du globe de Paris (doctoral grant from the French Ministère de l'Enseignement Supérieur de la Recherche et l'Innovation, ED STEP'UP 560), a doctoral grant from Sorbonne Université (ED GRNE) and ANR V-Care-18-CE03-0010.

VI.2.9. Author contributions

VI.2.10. Additional information

Competing interests: All the authors declare no competing interests.

CHAPTER VI.3: CHAOTIC MIXING EXPERIMENT CONSTRAINING CRYSTALLIZATION AND MIXING PROCESSES

Chaotic mixing experiment constraining crystallization and mixing processes

Lea Ostorero First draft that needs to be reviewed by our co-authors. imed journal after revisions of the draft: Earth and Planetary Science Letters.

Highlights:

• Crystallization during mixing unravelled by chaotic magma mixing experiment • Fast experimental growth rates of plagioclase • Basaltic end-member homogenized in a few hours • Magma mixing occurred in minutes to hours before the 2010 eruption of Kizimen

Abstract

The injection of a fresh primitive magma into a reservoir containing an evolved magma may promote magma mixing and consequently trigger explosive eruptions. The time needed to trigger eruptions is a fundamental question in modern volcanology but is still debated. The textures and chemical magma compositions produced by magma mixing have been used as chronometers to estimate the time interval between mixing processes and eruptions. During the last ten years, the kinetics and timescales of magma mixing have been investigated experimentally by performing chaotic mixing experiments. However, crystallization, which controls the chemical and textural evolution of magmas at pre-eruptive conditions

has not yet been considered together with magma mixing; the two processes have always been investigated separately. Here, we focus on the 2010 eruption of Kizimen (Kamchatka) as a representative case study of an explosive eruption triggered by magma mixing shortly prior to the eruption. This magma mixing/mingling event was preceded by a magma mixing event ~50 years ago that did not directly lead to the eruption of Kizimen. We perform for the first-time chemical mixing at sub-liquidus conditions (at which crystallization is promoted) via a chaotic mixing experiment in order to constrain the timescales of this first and old mixing event, and understand the pre-eruptive magmatic processes that followed. The experimental results achieved in this study indicate that the basaltic endmember was homogenized in a few hours, leading to hybrid compositions in the form of crystallized enclaves, and some mafic filaments. Experimentally generated plagioclase crystals show a fast growth by using the crystal size distribution. The triggering of the eruption by inhomogeneous magma mixing, occurred in minutes to hours before the 2010 eruption. This study provides great information for future monitoring of eruptions, knowing the time between the magmatic event and the eruption. The polished sections of the experimental and natural samples were analysed under Scanning Electron Microscopes (SEM): Zeiss Supra 55VP and the Carl Zeiss EVO MA10 SEM (Sorbonne Université, Institut des Sciences de la Terre de Paris, ISTeP, Paris and at the PARI platform at the Institut de physique du globe de Paris (Université Paris Cité), respectively) using an acceleration voltage of 15 kV and a beam current of 8 nA. The overview scanning of each experimental slice were composed of more than 1000 images (4 slices investigated). The SEM investigation is crucial for identifying the mixing textures, filaments (grey levels of different intensities) and crystals to be then chemically analyzed by the electron microprobe analyzer (EPMA) (70 pictures taken).

Thin section observations of the banded lavas were also carried out, to identify the contacts between the dacitic and andesitic end-members and to examine the crystals (60 images). VI.3.2.4.2. Electron Microprobe Analyzer (EPMA)

The experimental slices have been analysed for major elements (Si, Na, Ti, Al, Ca, K, Fe, Mg, Mn, P, Cl) with an acceleration voltage of 15 kV, a beam current of 4 nA and an unfocused beam of 6 µm to minimize alkali loss in the glass analysis (CAMECA SX-Five and SX-100; Service Camparis, Paris).

Internal standards used for glass measurements intercalibration between EPMA sessions were analysed in the same conditions as the residual glasses [START_REF] Balcone-Boissard | Non-volatile vs volatile behaviours of halogens during the AD 79 plinian eruption of Mt. Vesuvius, Italy[END_REF]. A total of 38 profiles through the glasses were obtained with a spacing between two points of 10-15 µm (Supplementary Table VI. 3.3). The same conditions were used to analyse the crystals but with a focused beam.

VI.3.2.5. Crystal Size Distribution (CSD)

The CSD theory is based on the principle that the distribution of crystals in a rock will depend on nucleation and growth rate, as well as the residence time in a magmatic chamber (or during an experiment) [START_REF] Cashman | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF][START_REF] Marsh | Crystal size distribution (CSD) in rocks and the kinetics and dynamics of crystallization[END_REF]. This technique allows us to have an estimation of the growth rate for the plagioclase in the experimental samples (knowing the duration of the

VI.3.6. Competing interests

The authors declare no competing interests.

VI.3.7. Author contributions

All authors contributed to the study conception and design. Material preparation and the experiment were performed by L. Ostorero and D. Morgavi. Data collection and analysis were performed by L. Ostorero. MELTS This basaltic end-member has been obtained by mixing oxides powders in the laboratory at the Department of Physics and Geology (University of Perugia) (Na2O and in K2O were converted in sodium and potassium carbonates (Na2CO3 and K2CO3); all FeO was converted in Fe2O3 and no MnO and P2O5 were added). After grinding with an agate mortar, a fine powder was obtained. VI.3.A.1.2

. Theoretical viscosities of the endmembers

Based on the starting materials compositions (Table VI. 3.1), we calculated the theoretical viscosity evolution of each end-member between 1100-1250°C [START_REF] Giordano | Viscosity of magmatic liquids: A model[END_REF]. We also show the evolution of the viscosity for the bulk synthetic basalt and dacite and the melts using the data calculated with MELTS software, with the effect of the presence of crystals [START_REF] Gualda | Rhyolite-MELTS: a Modified Calibration of MELTS Optimized for Silica-rich, Fluid-bearing Magmatic Systems[END_REF] (Supplementary Figure VI.3.2). VI.3.A.1.3. High-temperature magma mixing experiment

The COMMA experimental is presented in details in Morgavi et al. (2015) and the method described further down is from the same study. Two parameters define the COMMA geometry: (1) the ratio of the radii between the two cylinders (𝑟

) and ( 2) the eccentricity ratio to the outer cylinder (𝜀 = 𝛿 𝑅 𝑜𝑢𝑡 = 0.3), where δ is the distance between the centers of the inner and outer cylinders (Figure (Morgavi et al., 2015). Three main objectives are accomplished by the device: (a) allow independent rotations of the inner and outer cylinders, at variable speeds, directions and stirring protocols allowing to trigger chaotic mixing [START_REF] Swanson | A comparative computational and experimental study of chaotic mixing of viscous fluids[END_REF], (b) to put the apparatus in the hot spot region of a well-calibrated hightemperature furnace (up to 1800°C) and (c) finally to cool rapidly the silicate melts after the application of the mixing protocol to freeze in time the evolution of the mixing process, avoiding any crystallization of the sample. Two cylinders were drilled out of the crucible with the dacitic melt to have the required geometric configuration for the experiment: one cylinder for the spindle and a second cylinder for the initial position of the second end-member (synthetic basalt). The crucible (outer cylinder) is then connected to the motor above the furnace and was located off-center. Two alumina (Al2O3) rods are fixed to the upper and lower motor axis to act as the inner cylinder and holding the crucible (respectively). The lower end of the upper alumina rod is sheathed with a spindle (Pt80Rh20) to prevent from any contamination of the silicate melts during the experimental runs. Pt-PtRh thermocouples (accuracy ± 0.5 °C) are used to measure the temperature within the samples during the experiments (via the alumina rod). The vertical movement of the furnace allows to prepare the setup of the experiment by positioning the outer and inner cylinders outside the furnace. The furnace is then brought to the experimental temperature when the sample is still outside of the furnace (1200 °C). The furnace is then moved quickly in the end position required to do the experiment to minimize the time of attainment of thermal equilibrium (~10-15 min) in the sample. When the thermocouples inside the spindle measure the experimental temperature, the experiment is initiated by turning on the experimental mixing protocol.

VI.3.3c)

The computer-controlled mixing protocol consists of a combination of separated alternating rotations of outer and inner cylinders in this sequence: Supplementary Table VI To understand the dynamics of magmatic systems, one must first seek to characterize the time-dependent behavior of magma storage and ascent. Herein, we do this through a combination of the Crystal System Approach and careful study of Fe-Mg interdiffusion in orthopyroxene. This allows us to trace the pre-eruptive dynamics of magma plumbing systems, both in space and time. We apply this novel approach on two large silicic eruptions (about 3-5km 3 DRE/eruption) that occurred in the central part of Dominica Island (Lesser Antilles Arc): the eruptions of Layou (∼51 ka) from Morne Diablotins, and Roseau (∼33 ka) from Morne Trois Pitons-Micotrin. For the Roseau eruption, two magmatic environments (MEs) are identified on the basis of orthopyroxene composition, with a dominant reverse-zoning pattern from 50 to 54 to 54-59 mol% enstatite (En), indicating interaction with hotter magma. For the Layou eruption, three MEs are observed as represented by three populations of pyroxenes: En47-51, En51-53 and En53-58. The normal-zoning pathway from En51-53 to En47-51 is significantly registered by crystals, interpreted as convective mixing in a zoned reservoir. The reversezoning pathway from En47-51 to En51-53 and also En53-58 is also significantly present, supporting the mixing within the zoned reservoir but also suggesting mixing with a hotter magma, possibly stored in another part of a sub-volcanic mush. The crystal and glass compositions (melt inclusion and matrix glass) from both studied eruptions suggest heating and mixing between different magma pockets located within the mush that were the dominant process for mobilizing eruptible magma. In parallel, we constrain the associated pre-eruptive timescales by modeling the diffusive relaxation of Fe-Mg chemical gradients that originated within the zonation of the same orthopyroxene crystals. Diffusion modeling was considered along the b-axis of 66 zoned orthopyroxene crystals for these two eruptions, at a magmatic temperature of 850 ± 25 °C. In light of these results, we propose that the Layou and the Roseau magma reservoirs were rejuvenated and heated by ∼25-50 °C about 10 years prior to eruption by the injection of an underplating, hotter magma, creating the observed dominant reversezoning patterns of the erupted orthopyroxenes. We thus have evidence that silicic mush

Analytical Techniques

Textural Investigations

All crystals were investigated using Back-Scattered Electron (BSE) imaging via Scanning Electron Microscopy (SEM; ZEISS-Supra 55, Sorbonne Université, Paris) with an accelerating voltage of 20 kV, a beam current of 8 nA and a counting time in order to reduce noise impact on each image and thus better identify the crystal zonation (scan speed: 7; number of passages per scan: 8; Figure 2). The SEM images were processed using the free software ImageJ (http://rsb.info.nih.gov/ij/). Selected zones on orthopyroxene crystals have been imaged under high resolution to obtain a high spatial resolution profile (Figures 2B C).

Electron Probe Microanalyses

Major elements (Si, Ti, Al, Mn Fe, Mg, Ca, Na, K, P) in orthopyroxene crystals were analyzed by electron microprobe (Cameca SX-Five and SX-100, Service Camparis, Paris) with an accelerating voltage of 15 kV, a beam current of 10 nA and a focused beam. Counting time on peak and background was set at 10 s for all elements, except for Fe and Mg with 60 s for orthopyroxenes. Zoned orthopyroxene crystals have been analyzed along 2 µm step rim-to-rim or rim-to-core traverses (in regions previously picked out via high-resolution BSE imaging; Figure 2D).

Diffusion Modeling

The method (Figures 2A-E) is based on the modeling as a diffusion profile of the core-to-rim traverse (Figure 2E) obtained by coupling a high analytical resolution profile by Electron Probe Microanalyses (Figure 2D) and a high spatial resolution profile by SEM (Figure 2C), following the methodology developed by Allan et al. (2013) and Couperthwaite et al. (2020).

From each BSE image, a greyscale profile was extracted (Figure 2C), averaging over a selected area. In effect, this averages together many individual pixel traverses across the image, necessary to reduce thermal noise from the backscatter detector (Figure 2B). Profiles have been measured across the caxis, in the b-crystal direction and away from the terminal faces of crystals, which are more sensitive to recrystallization (Figure 2A). The modeling was accomplished using 1-D modeling, an acceptable procedure as the profiles are very short (∼20-30 µm) compared to the scale of the crystal (∼300 µm), and diffusion is effectively normal to the long face with no additional components. As the diffusion coefficient of Ganguly and Tazzoli (1994) is composition-dependent, modeled profiles develop asymmetry that is more pronounced as the compositional contrast becomes greater. The modeling of such profiles can be eased by exploiting the property that 1-D diffusion profiles proportionally share symmetry, so that if the approximatively at the same pressure/depth. Interestingly, there is only one direct transfer between ME L 1 and ME L 3, compared to significant transfers from ME L 1toME L 3vi aME L 2. This suggests that ME L 1 and ME L 3 may have been more isolated in space and/or in time (Figure 8D). Overall, the CSA depicts a stratified or cooling system, with limited exchanges, prior to a significant disruption of the cooler endmember by a slightly hotter magma, likely occurring within a decade before eruption (Figures 6,8; Table 2). In the CSA

SUPPLEMENTARY MATERIAL

The Supplementary Material for this article can be found online at: https://www.frontiersin.org/articles/10.3389/feart.2020.574032/ full#supplementary-material Conflict of Interest: The authors declare that the research was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

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