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Résumé en français

Il y a 2.4 Ma, le système climatique de la Terre est entré dans les Âges de Glace (Ice Ages, [START_REF] Imbire | Ice Ages -Solving the Mystery[END_REF], une période caractérisée par le développement et la retraite, quasi cycliques, de grandes calottes glaciaires continentales, rythmées par des variations de l'insolation sous le contrôle de variations subtiles des paramètres orbitaux de la Terre. Depuis les derniers ~12 ka, l'homme se trouve dans un état climatique dit interglaciaire, l'Holocène, caractérisé par des calottes glaciaires réduites (et donc un niveau de la mer élevé), des températures moyennes douces et une forte concentration de CO 2 atmosphérique [START_REF] Bereiter | Revision of the EPICA Dome C CO2 record from 800 to 600 kyr before present[END_REF]. Si le climat de la Terre suivait toujours un modèle naturel, glaciaire-interglaciaire, la Terre entrerait progressivement dans une nouvelle période glaciaire à un moment donné dans l'avenir en réponse aux changements dans la distribution latitudinale et saisonnière de l'insolation et aux mécanismes de rétroaction qui déterminent la concentration atmosphérique en CO 2 . Cependant, le climat de la Terre ne suit plus un chemin naturel. Les activités humaines, associées à la révolution industrielle et l'utilisation sans cesse croissante du carbone fossile, ont porté le niveau de concentration du CO 2 atmosphérique (et d'autres gaz à effet de serre) à un niveau sans précédent au cours des 800 000 dernières années [START_REF] Bernstein | climate change 2007: synthesis report[END_REF]. Il est maintenant clair que le climat de la Terre connaît un changement global, qui est induit par l'Homme [START_REF] Bernstein | climate change 2007: synthesis report[END_REF]. Comment le climat mondial continuera-t-il à changer à l'avenir sous l'influence des activités humaines? C'est une question cruciale pour les êtres humains et l'ensemble des êtres vivants sur Terre. Les modélisateurs du climat ont la responsabilité de prévoir cette évolution en développant des modèles numériques toujours plus sophistiqués. Mais le système climatique de la Terre est extrêmement complexe, avec de nombreuses boucles de rétroaction et des mécanismes clés qui sont encore mal compris et ne peuvent pas encore être modélisés correctement. Pour prédire avec précision les changements climatiques futurs, il faut une bien meilleure compréhension des interactions complexes au sein du système climatique, qui impliquent des couplages étroits entre la biosphère, l'hydrosphère, la cryosphère et l'atmosphère.

Si l'on veut s'attaquer à la complexité des mécanismes climatiques de la Terre à différentes échelles de temps qui sont pertinentes pour l'évolution à court et à long terme de l'humanité et des sociétés humaines, le passé détient la clé de l'avenir. Au cours des dernières décennies, changements du climat passé ont été reconstitués à partir de diverses archives naturelles (p. ex. carottes de glace, sédiments océaniques et stalagmites des cavernes, cernes d'arbres, etc.) Sur la base de ces données, les paléoclimatologues ont pu montrer que la Terre a connu des changements climatiques importants -et parfois brusques -dans le passé. Les plus importants de ces changements sont dus aux variations de l'insolation associées à l'orbite terrestre et aux changements de la pression partielle en CO 2 atmosphérique, et/ou résultent de l'instabilité de la calotte glaciaire et des modifications de la redistribution de la chaleur et du sel dans l'océan.

Bien que nous ayons acquis une connaissance toujours plus fine de la variabilité du climat de la Terre, il y a encore beaucoup d'informations à apprendre de l'étude des changements Par exemple, au cours de la dernière période glaciaire, une série d'épisodes de froid (stades) ont eu lieu dans l'Atlantique Nord (Barker et al., 2011). Ils étaient dus à une réduction du transport de chaleur interhémisphérique, vers le Nord, en réponse à une circulation méridionale atlantique (AMOC) plus léthargique (Böhm et al., 2015). Au cours de la même période, l'hémisphère sud a connu des épisodes plus chauds en raison de l'accumulation océanique de chaleur qui n'était plus transportée efficacement vers l'hémisphère nord (Buizert et al., 2018). Ce mécanisme de refroidissement dans un hémisphère et de réchauffement dans l'autre est connu sous le nom de mode de bascule bipolaire [START_REF] Stocker | A minimum thermodynamic model for the bipolar seesaw[END_REF]. Cette asymétrie de température interhémisphérique dans les océans affecte la circulation atmosphérique à l'échelle planétaire en influant sur la température et, par conséquent, sur les gradients de pression atmosphérique.

Elle a un impact sur la zone de convergence intertropicale (ITCZ), la ceinture pluviométrique la plus intense de la Terre, qui tend à migrer vers l'hémisphère le plus chaud (Schneider et al., 2014). La migration de la position moyenne de la ITCZ exerce un effet substantiel sur le cycle hydrologique des zones tropicales appartenant à la ceinture mondiale de mousson (par exemple, Carolin et al, 2013;Cheng et al, 2016;Gibbons et al, 2014 ;Levi et al, 2007).

Mais les impacts climatiques touchent aussi d'autres parties du globe puisque l'atmosphère et l'océan sont étroitement liés et exercent des influences mutuelles l'un sur l'autre, contribuant à propager les changements à distance. Par exemple, au cours de la dernière période glaciaire, le déplacement vers le Sud de la ITCZ pendant les épisodes climatiques froids (chauds) dans l'hémisphère nord (sud) a été associé aux déplacements vers le Sud de la ceinture des vents du Sud-Ouest (SSW) [START_REF] Anderson | Wind-Driven Upwelling in the Southern Ocean and the Deglacial Rise in Atmospheric CO2[END_REF], permettant au front subtropical de l'Océan Austral (STF) de migrer vers l'Antarctique. Cette situation a permis la fuite d'un plus grand flux d'eaux chaudes et salines dans l'Océan Atlantique via le courant des Aiguilles, autour de la pointe sud de l'Afrique [START_REF] Bard | Migration of the subtropical front as a modulator of glacial climate[END_REF]Beal et al, 2011 ;Peeters et al, 2004). L'augmentation du flux d'eau par le passage des Aiguilles -grâce au renforcement de l'AMOC (c'est-à-dire le pompage de plus de chaleur dans l'Océan Atlantique nord et le réchauffement de l'hémisphère nord) -a réduit le contraste thermique interhémisphérique, ramenant le SSW et le STF à leurs positions la plus au Nord. Ce changement de position en réduisant ou en stoppant le passage des eaux du courant des Aiguilles vers l'Atlantique, s'est traduit par un affaiblissement de l'AMOC qui a permis de refroidir à nouveau l'hémisphère nord (Beal et al., 2011). Le système que nous venons de décrire est donc une boucle de rétroaction complexe. Les changements du gradient thermique interhémisphérique peuvent forcer la réorganisation de la circulation atmosphérique et océanique. Ensuite, la modification de la circulation atmosphérique et océanique influe sur la redistribution de la chaleur et de l'humidité à l'échelle planétaire, ce qui, à son tour, influe sur les différences de température entre les hémisphères.

Comme nous venons de le voir, les tropiques et les océans tropicaux sont au coeur d'interactions complexes et interhémisphériques. Les océans tropicaux ne jouent pas seulement un rôle de transition entre les deux hémisphères. Principaux collecteurs de la chaleur apportée par rayonnement solaire, sources d'une évaporation active et de puissants processus convectifs, les océans tropicaux constituent une source majeure de chaleur (directe ou latente) et de vapeur d'eau pour le climat mondial. L'Indo-Pacific Warm Pool (IPWP), la zone océanique la plus chaude de la Terre, a été appelé le « moteur thermique à vapeur » du globe [START_REF] De Deckker | The Indo-Pacific Warm Pool: critical to world oceanography and world climate[END_REF].

La position moyenne de l'ITCZ-, qui est sensible au gradient thermique inter-hémisphérique, influe sur la variabilité hydrologique dans la région de l'IPWP (Gibbons et al., 2014). Le transfert d'eau qui s'effectue via le courant des Aiguilles est lui-même alimenté par les eaux en provenance de l'IPWP, transportées dans le cadre du flux d'eau indonésien (Indonesian Thourghflow, ITF), un élément clé de la redistribution de la chaleur -entre le Pacifique et l'océan Indien. Par conséquent, l'IPWP est un domaine particulièrement important en ce qui concerne les changements climatiques à grande échelle et l'étude de son évolution dans le passé est essentielle pour mieux comprendre les mécanismes qui déterminent le climat mondial.

L'objectif principal de cette thèse est (i) de reconstruire le changement hydroclimatique de l'IPWP et de discerner les mécanismes de télé-connexion couplant les changements climatiques à basse et haute latitude et leur impact sur les climats à basse-latitude (ITCZ) ; et (ii) de reconstruire la variabilité de l'ITF au cours des derniers cycles climatiques, à des échelles de temps allant des variations orbitales aux changements millénaires, afin de mieux comprendre les processus qui contrôlent les transports de chaleur et d'eau douce du Pacifique vers l'Océan Indien. Ainsi, en un sens, ce travail de thèse porte sur l'interaction des principales caractéristiques des circulations méridionale et zonale qui caractérisent l'IPWP.

Afin d'atteindre ces objectifs, j'ai analysé deux carottes de sédiments situées dans l'IPWP, à la sortie de l'ITF dans l'Océan Indien : la carotte MD98-2165 (9,65°S, 118,34°E, profondeur d'eau de 2100m) et la carotte MD98-2166 (11,96°S, 120,46°E, profondeur d'eau de 3875m). J'ai combiné mes résultats aux données publiées d'autres carottes sédimentaires dans l'IPWP ou l'Océan Indien. Les données paléoclimatiques/paléocéanographiques seront principalement obtenues à partir de la carotte la moins profonde (MD98-2165), dont les tests foraminifères ne sont pas fortement affectés par la dissolution. La carotte la plus profonde, MD98-2166, a été utilisée exclusivement, en conjonction avec le MD98-2165, pour mieux estimer l'impact de la dissolution sur le contenu en éléments traces des tests de foraminifères.

La température de la surface de la mer (SST) et la température de l'eau de la thermocline (TWT) ont été reconstituées par la paléothermométrie Mg/Ca sur des coquilles de foraminifères. La salinité de surface de la mer (SSS) a été évaluée en reconstruisant le δ18O d'eau de mer par la combinaison de Mg/Ca-T° et du δ18O de tests de foraminifères planctoniques.

Avant d'utiliser la paléothermométrie Mg/Ca, l'impact des méthodes de nettoyage (avec ou sans étape réductrice pour éliminer les oxydes) a été évalué via des analyses de foraminifères planctoniques foraminifères Mg/Ca par ICP-MS. Ce travail est obligatoire si l'on veut combiner avec précision les températures estimées à partir du traceur Mg/Ca lors d'études ayant utilisé différentes méthodes de nettoyage dans la région de l'IPWP.

Ce mémoire est divisé en six chapitres. Trois de ces chapitres (3, 4 et 5) ont été rédigés sous forme de manuscrits autonomes qui seront publiés dans des revues à comité de lecture. En raison de l'architecture particulière de cette thèse, certaines redondances sont inévitables, notamment en ce qui concerne le matériel, les méthodes et les cadres modernes. En raison du format court des articles scientifiques, de nombreux détails n'ont pas pu être inclus dans les chapitres qui seront soumis pour publication et nous avons donc décidé d'écrire un chapitre spécifique "Cadre moderne et passé" et un chapitre "Matériel et méthode" pour donner un aperçu plus complet du domaine d'étude, des carottes utiliséss et des méthodes (incluant de nombreux résultats analytiques non exploités dans les chapitres principaux). Chapitre 1. Ce chapitre traite de la climatologie et de l'océanographie modernes et passées de la région étudiée, en mettant l'accent sur l'IPWP, l'ITCZ et l'ITF.

Le chapitre 2 se concentre sur le matériel et les méthodes. Il résume et complète les éléments d'information disponibles dans les chapitres spécifiques construits comme des manuscrits autonomes pour soumission dans des revues à comité de lecture. Il fournit aux lecteurs des éléments supplémentaires sur les carottes MD98-2165 et MD98-2166 et approfondit certains aspects méthodologiques qui seront abordés superficiellement dans les chapitres. Dans ce chapitre, je mettrai également l'accent sur la reconstruction des propriétés hydrologiques passées (c.-à-d. SST, SSS et TWT).

Chapitre 3. Ce chapitre se concentre sur l'effet des méthodes de nettoyage sur l'analyse Mg/Ca dans trois espèces de foraminifères planctoniques -Globigerinoides ruber, Pulleniatina obliquiloculata et Globorotalia menardii. Nous avons évalué les effets de deux méthodes de nettoyage (avec ou sans étape réductrice) sur la reconstruction des températures de l'eau à partir du traceur Mg/Ca. Ce chapitre met l'accent sur la dépendance à l'espèce et le rôle de l'état de préservation des foraminifères en ce qui concerne les impacts du nettoyage sur la Mg/Ca. Nous montrons également l'impact de l'offset lié au nettoyage et proposons une stratégie simple pour corriger cet impact afin de construire des bases de données Mg/Ca-T° plus cohérentes et robustes pour les reconstructions océanographiques et climatiques passées. Plusieurs autres éléments (à savoir : Li, B, Sr, Ba, U, Mn, Mn, Fe et Al) ont été mesurés simultanément avec le Mg. Les ratios de ces éléments par rapport au calcium sont fournis en chiffres à la fin de la partie annexe de la thèse.

Le chapitre 4 se concentre sur les variations hydrologiques dans la région IPWP et sur les téléconnexions des changements climatiques à basse et haute latitude par la circulation atmosphérique. Dans ce chapitre, nous présentons une reconstruction de 270 000 ans de la composition de l'eau de mer locale (δ 18 O sw ), un indicateur de la salinité et de l'hydrologie de surface de la mer, estimé en combinant Mg/Ca-SST et δ 18 O de G. ruber dans la carotte MD98-2165. Ce chapitre traite plus particulièrement de la relation entre les changements hydrologiques dans l'IPWP et les variations passées du gradient thermique inter-hémisphérique et de la dynamique de l'AMOC.

Chapitre 5. Ce chapitre se concentre sur la variabilité du transport ITF au cours des deux derniers cycles climatiques. A cette fin, nous analysons les changements passés de la température de la thermocline à la sortie de l'ITF dans l'Océan Indien (dérivés du Mg/Ca-T° de l'espère planktonique profonde P. obliquiloculata dans la carotte MD98-2165) et nous comparons nos résultats avec divers enregistrements de TWT de l'archipel Indonésien, dont un enregistrement récent situé dans l'upwelling situé à l'Ouest de Sumatra. Ces comparaisons permettent de mieux déchiffrer l'évolution de l'intensité de l'ITF dans le contexte de la forte dynamique de la couche supérieure de l'océan due aux vents qui affectent la majeure partie de l'archipel indonésien.

Le mémoire se termine par un bref chapitre de conclusion et de perspective (chapitre 6).

Les principales conclusions de ce travail sont :

(1) Dans l'analyse des foraminifères Mg/Ca, l'introduction d'une étape « réductrice » dans le nettoyage des foraminifères entraîne une perte moyenne de Mg/Ca d'environ 6 ± 7% (1s, n=328) lorsqu'on combine les résultats des trois espèces G. ruber, P. obliquiloculata et G. menardii du noyau MD98-2165 et MD98-2166.) Cette perte équivaut à une diminution de 0,7 ± 0,9°C des températures reconstruites. Cette perte résulte de la dissolution partielle induite par l'étape de nettoyage réducteur. Mais nos données indiquent également que les différences de Mg/Ca induites par le nettoyage dépendent de l'espèce de foraminifère et reflètent également l'état de conservation des tests. Cela explique pourquoi certaines incohérences ont été observées dans la littérature concernant les effets du nettoyage sur le Mg/Ca. Bien que nous ayons mis en évidence des différences claires de Mg/Ca induites par le nettoyage, il n'y a aucun signe de changement significatif dans les variations relatives de Mg/Ca le long d'une carotte donnée et pour une espèce donnée lorsque l'on compare les courbes obtenues par les deux méthodes de nettoyage.

En d'autres termes, les changements relatifs de température reconstitués le long des enregistrements sédimentaires (par exemple, le changement de température entre le LGM et l'Holocène) ne sont pas affectés par le choix de la méthode de nettoyage. Cependant, comme les températures absolues de l'eau de mer sont biaisées, la communauté paléocéanographique doit développer des stratégies claires pour construire des températures cohérentes basées sur le traceur Mg/Ca. Nous suggérons d'utiliser de préférence l'approche de nettoyage non réductrice pour construire des ensembles de données homogènes de températures. L'étape de réduction ne doit être incluse que lorsqu'une couche d'oxyde importante est observée (et, ensuite, les températures Mg/Ca résultantes doivent être corrigées pour l'offset créé par le nettoyage réducteur, en utilisant soit un facteur de correction global (c.-à-d. +6%) ou, mieux, en appliquant des facteurs de correction spécifiques à chaque espèce. Bien que nous ayons clairement établi que la dissolution in situ a un impact sur le décalage Mg/Ca lié au nettoyage, nous sommes conscients que la dissolution est un mécanisme complexe et que nos deux carottes ne nous ont donné qu'un aperçu partiel de l'impact de la dissolution. D'autres carottes situées à différentes profondeurs d'eau dans les mêmes zones seront nécessaires pour explorer pleinement l'impact de la dissolution sur les biais de nettoyage du Mg/Ca.

(2) Nous avons montrons qu'il existe de faibles corrélations et parfois des relations presque antiphasées entre le δ 18 O sw-ivc de la carotte MD98-2165 et le δ 18 O d'une stalagmite de l'île voisine de Flores, indiquant clairement que les changements passés de salinité dans la mer de Timor ne sont pas principalement contrôlés par les changements du régime local des précipitations. Au lieu de cela, nous avons montré que le δ 18 O sw-ivc de la carotte MD98-2165 variait de façon cohérente avec un enregistrement δ 18 O d'une stalagmite de Bornéo, et que les deux enregistrements sont fortement corrélés également à des changements d'échelle millénaire au signal δ 18 O (préalablement corrigé d'un fort signal de préession) d'une stalagmite chinoise. Nos résultats soulignent donc (i) l'existence de modes cohérents de changements hydroclimatiques enregistrés sur la majeure partie de la région de mousson australo-asiatique, et (ii) le fait que les eaux de surface sortant de l'Archipel Indonésien dans le cadre de l'ITF transportent un signal δ 18 O sw-ivc qui est cohérent avec les enregistrements de l'hémisphère nord (Bornéo, Chine) et parfois en anti-phase avec des changements locaux, des précipitations de l'hémisphère sud (Flores). Tous ces éléments peuvent s'expliquer par les changements passés dans la dynamique de la ITCZ. De grandes excursions positives de l'enregistrement MD98-2165 δ 18 O sw-ivc se sont produites à chacune des trois dernières terminaisons glaciaires et à la transition entre MIS 4 et MIS3. Ces changements correspondent avec des périodes de ralentissement de l'AMOC comme le révèle un enregistrement Pa/Th récent. Cela suggère que les changements rapides de l'hydroclimat tropical oriental de l'Océan Indien sont affectés par le forçage de l'hémisphère nord via la dynamique de l'ITCZ. Les oscillations en phase enregistrées dans l'enregistrement δ 18 O sw-ivc de la carotte MD98-2165, l'enregistrement δ18O de la stalagmite de la grotte de Bornéo et l'enregistrement ∆δ 18 O de la stalagmite de la grotte de Chine indiquent que les variations du gradient thermique interhémisphérique ont toujours joué un rôle majeur dans les changements hydroclimatiques tropicaux aux échelles millénaires à orbital.

(3) Nous avons montré que les changements dans les TWT dérivés de P. obliquiloculata Mg/Ca sur le site MD98-2165 reflètent l'équilibre entre la force de l'ITF et l'intensité des vents de mousson locaux du sud-est (SE) qui entrainent un phénomère d'upwelling. Aux échelles de temps millénaire à orbitale, des vents de mousson du sud-est plus forts entraînent un renforcement du transport ITF. A l'échelle glaciaire-interglaciaire, l'ITF est modulé par l'évolution du niveau de la mer, avec un affaiblissement de l'ITF lié à la baisse du niveau de la mer. Nos résultats montrent que l'ITF était généralement faible pendant les phases froides MIS 8, MIS 6 et MIS 2-4, alors que l'ITF était plus fort MIS 7, MIS 5 et pendant MIS 1. En particulier, le transport de chaleur par l'ITF était maximum aux cours des périodes marquées par une insolation estivale maximale dans l'hémisphère nord.

General introduction

2.4 Ma ago, the Earth's climate system entered the Ice Ages [START_REF] Imbire | Ice Ages -Solving the Mystery[END_REF], a period characterized by the quasi-cyclic waxing and waning of large polar ice caps, paced by changes in insolation under the control of subtle variations of Earth's orbital parameters. Since the last ~12 ka, human beings have been leaving in a so-called interglacial climate state, the Holocene, characterized by reduced ice caps (and thus a high sea level), mild average temperatures and a high level of atmospheric CO 2 [START_REF] Bereiter | Revision of the EPICA Dome C CO2 record from 800 to 600 kyr before present[END_REF]. If the Earth's climate would be still following a natural, glacial-interglacial pattern, the Earth would enter progressively into a new glacial period at some point in the future in response to changes in the latitudinal and seasonal distribution of insolation and to feedback mechanisms driving the atmospheric pCO 2 level.

However, the Earth's climate is not following a natural path any longer. Human activities, boosted by industrial revolution and the ever-growing use of fossil carbon, have increased the atmospheric pCO 2 (and other greenhouse gas) to a level unprecedented during the past 800,000 years [START_REF] Bernstein | climate change 2007: synthesis report[END_REF]. It is now clear that the Earth climate is experiencing a global change, which is man-induced [START_REF] Bernstein | climate change 2007: synthesis report[END_REF]. How will the global climate continue to change in the future under the influence of human activities is a crucial issue for human beings and living creatures on Earth. Climate modelers have the responsibility to foresee this evolution through the development of ever more sophisticated numerical models. But the Earth's climatic system is extremely complex with many feedback loops and key mechanisms that are still poorly understood and cannot be modeled properly yet. To predict accurately future climate changes requires a much better understanding of complex interactions within the climate system, which involve tight couplings between the biosphere, the hydrosphere, the cryosphere and the atmosphere.

If one wants to tackle the complexity of Earth climate mechanisms at the different timescales that are relevant to short and long-term evolution of mankind and human societies, the Past holds the key to the Future. During the past decades, numerous paleoclimate records have been reconstructed based on various natural archives (e.g., ice cores, ocean sediments and cave stalagmites, tree rings, etc.). Based on these data, paleoclimatologists were able to show that the Earth has experienced important -and sometime abrupt -climate changes in the past. The most important of these changes were driven by variations in orbitally-driven insolation and atmospheric pCO2, and/or resulted from ice cap instabilities and modifications of the general redistribution pattern of heat and salt in the ocean.

Although we have acquired a growing knowledge on Earth's climate variability, there are still many more pieces of information to learn from the study of past climate changes. Such a knowledge is crucial to provide better constrains on climate change and anticipate its future evolution under a different radiative forcing. In the present work, I will more particularly address the role and changes of low-latitude mechanisms associated to the hydrologic cycle and the transfer of heat and salt through the Global Conveyor Belt.

With the increasing number of paleoclimate observations and the better understanding of today's climate, a global picture has emerged, which highlight the importance of atmospheric and oceanic interconnections. The heat and moisture redistributions by atmospheric and oceanic circulation patterns are seen as a key mechanism to understand and anticipate global climate evolution and its regional or local impacts. This redistribution is affected by many complex feedback loops and remote interactions.

For example, during the last glacial period, a series of cold episodes (stadials) occurred in the North Atlantic (Barker et al., 2011). They were due to the reduced, inter-hemispheric, northward transport of heat in response to a more sluggish Atlantic meridional overturning circulation (AMOC; (Böhm et al., 2015)). During the same period, the Southern Hemisphere experienced warmer episodes due to the oceanic accumulation of heat that was no longer transported efficiently to the Northern Hemisphere (Buizert et al., 2018). This mechanism of cooling in one hemisphere with warming in the other is known as the bipolar seesaw mode [START_REF] Stocker | A minimum thermodynamic model for the bipolar seesaw[END_REF]. This inter-hemispheric temperature asymmetry in the oceans affects the atmospheric circulation on a global scale by impacting temperature and, thus, atmospheric pressure gradients. It has an impact on the intertropical convergence zone (ITCZ), the most intense rainfall belt on Earth, which tends to migrate towards the warmer hemisphere (Schneider et al., 2014). The migration of the mean position of ITCZ exerts a substantial effect on the hydrologic cycle of tropical areas belonging to the global monsoon belt (e.g., Carolin et al., 2013;Cheng et al., 2016;Gibbons et al., 2014;Levi et al., 2007). But climate impacts reach other parts of the globe since the atmosphere and the ocean are tightly coupled and exert mutual influences on one another, helping to propagate changes remotely. For example, during the last glacial period, the southern shift of the ITCZ during cold (warm) episodes in the northern (southern) hemisphere, were associated to southward shifts of southern westerly winds (SSW) [START_REF] Anderson | Wind-Driven Upwelling in the Southern Ocean and the Deglacial Rise in Atmospheric CO2[END_REF], allowing the Southern Ocean subtropical front (STF) to migrate towards Antarctica, which made it possible the leakage of more warm and saline waters into the Atlantic Ocean through the Agulhas leakage zone, around the southern tip of Africa [START_REF] Bard | Migration of the subtropical front as a modulator of glacial climate[END_REF]Beal et al., 2011;Peeters et al., 2004). Such movements of circulation patterns were at the heart of flip-flop modifications of the climate. The increase in Agulhas leakage -though the strengthening of the AMOC (i.e. pumping more heat to the north Atlantic Ocean and warming the Northern Hemisphere) -reduced the inter-hemispheric thermal contrast, bringing the SSW and STF back to a northernmost position. This shift in position would reduce or suppress the Agulhas leakage, weaken the AMOC and cool the North Hemisphere again (Beal et al., 2011). The system we just described is a complex, feedback loop. Changes in the inter-hemispheric thermal gradient can force the reorganization of atmospheric and oceanic circulation. Then, the modified pattern of atmospheric and oceanic circulation influences the redistribution of heat and moisture globally, which in turn affects the temperature differences between the hemispheres.

As we have seen above, the tropics and tropical oceans are the heart of complex, interhemispheric interactions. The tropical oceans do not only play a transition role between the two hemispheres. As the main collectors of heat brought by direct insolation forcing and active evaporation and convective processes, the tropical oceans act as a major source of heat (direct or latent) and vapor for the global climate. The Indo-Pacific Warm Pool (IPWP), the warmest oceanic area on Earth, has been called the "heat" and "steam" engine of the globe [START_REF] De Deckker | The Indo-Pacific Warm Pool: critical to world oceanography and world climate[END_REF]. The mean position of overlying ITCZ, which is sensitive to the interhemispheric thermal gradient, can be constrained by the hydrologic variability in IPWP region (Gibbons et al., 2014).

The Agulhas leakage is fueled by waters from IPWP, which are carried as part of the Indonesian Throughflow (ITF), a key element in the heat redistribution by global ocean circulation between the Pacific and the Indian Ocean and from low to high latitude. Therefore, the IPWP is particularly important area as far as large scale climate changes are concerned and the study of its evolution in the past is pivotal to better understand the underlying mechanisms that drive global climate.

The main objective of this thesis is (i) to reconstruct the hydroclimate change in IPWP and discern teleconnection mechanisms between the low-and high-latitude climate changes and their impact on the low latitude climates (i.e., ITCZ); (ii) to reconstruct the variability of ITF at different timescales, from orbital to millennial, and discern the processes that control the ocean heat and fresh water transport from the Pacific to the Indian Ocean. Thus, in a sense, this PhD work addresses the interplay of the main meridional and zonal circulation features that characterize the IPWP.

In order to achieve these goals, I analyzed two sediment cores located in the IPWP, at the exit of the ITF into the Indian Ocean: core MD98-2165 and core MD98-2166 (11.96° S, 120.46° E, water depth of 3875m) and combined our results with published data from other sedimentary cores in the IPWP or the Indian Ocean.

Paleoclimate/paleoceanographic reconstruction will be gained chiefly from the shallower core MD98-2165, whose foraminiferal tests are not strongly affected by dissolution. The deeper core MD98-2166 has been exclusively used, in conjunction with MD98-2165, to better estimate the impact of dissolution on trace elements measured on foraminiferal tests. This dissertation is divided into six chapters. Three of these chapters (3, 4 and 5) have been written as stand-alone manuscripts to be published in peer-reviewed journals.

Because of the paper-ready architecture of this dissertation, some redundancies are inevitable, in particular concerning material, methods and modern settings. On the opposite, due to the short format of scientific papers, many details could not be included in the chapters that will be submitted for publication and we decided, therefore, to write a specific "Modern and Past settings" chapter, and a "Material and Method" chapter to provide a more complete overview of the study area, the cores used and the methods (including many analytical results not exploited in the chapters).

Chapter 1. This chapter deals with modern and past climatology and oceanography of the studied region, with specific highlights on the IPWP, the ITCZ and the ITF.

Chapter 2 focuses on the material and methods. It summarizes and completes the pieces of information available in the specific chapters built as stand-alone manuscripts for submission into peer-review journals). It provides the readers with additional items about cores MD98-2165 and MD98-2166 and digs more deeply into some methodological aspects that will be superficially discussed in the chapters. In this chapter, I will put also the emphasis on the reconstruction of past water properties (i.e., SST, SSS and TWT).

Chapter 3. This chapter focuses on cleaning methods effect on Mg/Ca analysis in three planktonic foraminifer species --Globigerinoides ruber, Pulleniatina obliquiloculata and Globorotalia menardii. we evaluated the effects of two cleaning methods (with or without a reductive step) on Mg/Ca-based water temperature reconstruction. This chapter puts the emphasis on the species-dependency and the role of preservation state of foraminifers regarding cleaning impacts on Mg/Ca. We also show the impact of cleaning-related offset and propose a simple strategy to correct for this impact in order to build more coherent and robust Mg/Ca-T° databases for past oceanographic and climatic reconstructions. Several other elements (namely: Li, B, Sr, Ba, U, Mn, Fe, and Al) were simultaneously measured with Mg. The ratios of these elements with respect to calcium are provided in figures in the end of thesis appendix part.

Chapter 4 focuses on the hydrologic variations in the IPWP region and the low to high latitude teleconnections of climate changes through atmospheric circulation. In this chapter, we present a 270, 000-year record of local seawater δ 18 O composition-a sea surface salinity and hydrology indicator-estimated by combining Mg/Ca-SST and δ 18 O of G. ruber in core MD98-2165. This chapter addresses more particularly the relationship of hydrologic changes in the IPWP with respect to past variations of interhemispheric thermal gradient and AMOC dynamics.

Chapter 5. This chapter focuses on the variability of ITF transport over the last two climatic cycles. For this purpose, we analyze past changes of the thermocline water temperature in the exit of ITF into the Indian Ocean (derived from P. obliquiloculata Mg/Ca-TWT in core MD98-2165) and compare our results with various TWT records from the Indonesian Archipelago, including a recent record located in the main upwelling located to the West of Sumatra. These comparisons make it possible to better decipher the relative evolution of ITF strength in the context of the strong wind-driven upper ocean dynamics that affect most of the Indonesian Archipelago.

The dissertation ends up with a short conclusion and perspective chapter (Chapter 6)

Chapter 1 : Oceanographic and climatic settings of the studied area

In this chapter, I present modern and past climatology and oceanography of the studied region, with specific highlights on the IPWP, the ITCZ and the ITF.

The Indo-Pacific Warm Pool

At present

The Indo-Pacific Warm Pool (IPWP) is the warmest oceanic area on Earth, with peak surface temperatures around ~ 34°C and an annual mean surface temperature above 28 °C (Figure 1.1).

The IPWP covers a surface area of about ~ 30 × 10 6 km 2 . It collects its energy both from the piling up of warm surface waters pushed westward by the trace winds and by direct insolation.

Due to heavy precipitations, moderate winds and net surface heating, the mixed layer is relatively shallow over the IPWP, with a mean depth of about 30-40 m. Warm temperatures extend down to about 200m, but the heart of the IPWP is shallow in nature with temperatures dropping below 25 °C within less than 100 m water depth. Although this warm and fresh thin skin of buoyant surface water occupies only 0.05% of the total ocean-water mass, it plays a leading role in driving atmospheric circulation due to the huge heat it contains, making it the most important energy source for the global circulation system [START_REF] De Deckker | The Indo-Pacific Warm Pool: critical to world oceanography and world climate[END_REF]. Small changes in tropical SST can lead to a marked increase in surface-ocean evaporation and water vapour in the atmosphere [START_REF] Flohn | Recent changes of the tropical water and energy budget and of midlatitude circulations[END_REF][START_REF] Webster | The role of hydrological processes in ocean-atmosphere interactions[END_REF]. Thus, any change in the temperature, size, or positioning of the Warm Pool in the past would have had a profound effect on global climate [START_REF] Clement | Orbital controls on the El Niño/Southern Oscillation and the tropical climate[END_REF].

The central part of IPWP around the equator is constantly warm over the whole year with small seasonal change (Figure 1.2). By contrast, the boundary of IPWP (defined by the 28 °C isotherm) vary significantly along with season changes. For example, from the austral summer (January) to the austral winter (July), the 28 °C isotherm in the southern Hemisphere moves from ~ 20°S to ~10°S. Similarly, in the northern Hemisphere part, the 28 °C isotherm can move between 25°N and 10°N on a seasonal basis. The thermocline water temperature (at a water depth of ~ 100m) is generally higher in the tropical western Pacific Ocean than the tropical eastern Indian Ocean and the Indonesian seas (Figure 1.3), indicating much deeper thermocline depth in the western Pacific sector. Seasonally reversing winds generate upwelling and dowelling along the coast of many islands in the Indonesian Archipelago region, leading to varying thermocline temperatures [START_REF] Nagai | Internal tides and associated vertical mixing in the Indonesian Archipelago[END_REF]Susanto et al., 2001).

Due to heavy evaporation and convective activity, surface water in the IPWP is rather fresh, the annual mean surface salinity being less than 35‰ (Figure 1.4). There are seasonal changes in surface salinity owing to the fact that the IPWP spreads over a region which is affected by a complex monsoon system connecting the South-East Asia and Australia: The Malaysian-Australian monsoon system. As we will see later in this chapter, monsoon dynamics and the seasonal shift of precipitation centers are directly related to the spatial movement of the Inter-Tropical Converge Zone (ITCZ). During the 'wet' season, salinity is affected by direct precipitation over the ocean as well as by freshwater addition from rivers, explaining why higher salinity changes occur in marginal seas. Away from the equator, huge pours of fresh water are associated to cyclonic activity. As a result of such 'dilution' in the Warm Pool region, the salinity of waters that are being transferred from the western Pacific Ocean into the eastern Indian Ocean will change, explaining why this region is sometime called the 'dilution basin' of the Great Conveyor Belt [START_REF] De Deckker | The Indo-Pacific Warm Pool: critical to world oceanography and world climate[END_REF].

In the past

Based on the study of ocean sediment cores, paleo-SST has been reconstructed from different sites in the IPWP region. Linsley et al., (2010) show that from the early to the middle Holocene (from 10 to 7 thousand years ago; "ka" hereafter), SST in the western Pacific sector of IPWP had been about 0.5 °C higher than during the pre-industrial times, suggesting a warmer and more extended IPWP. During the last glacial maximum (LGM, ~19 ka), however, SST of IPWP was significantly cooler, by 3 -4°C, compared to the early Holocene (e.g. Bolliet et al., 2011;Fan et al., 2018;[START_REF] Lea | Climate impact of late quaternary equatorial Pacific sea surface temperature variations[END_REF]Levi et al., 2007;[START_REF] Mohtadi | Palaeoclimatic insights into forcing and response of monsoon rainfall[END_REF]Visser et al., 2003;Xu et al., 2008). This means that, if one strictly applies the modern IPWP definition (i.e. a region over which annual mean SST is above 28°C), the IPWP was gone during the LGM. Of course, this is not the case. The IPWP was still the warmer area of the world ocean, although the rest of the glacial oceans were also colder at that time [START_REF] Lea | Climate impact of late quaternary equatorial Pacific sea surface temperature variations[END_REF].

What drove SST change since the LGM in the IPWP region is still unclear. Atmosphere pCO 2 is thought to have played a critical role in forcing SST change in western Pacific sector of IPWP [START_REF] Tachikawa | CO2 radiative forcing and Intertropical Convergence Zone influences on western Pacific warm pool climate over the past 400ka[END_REF]. [START_REF] Koutavas | Temperature correlations between the eastern equatorial Pacific and Antarctica over the past 230,000 years[END_REF] argues that the tropical SST change since the LGM was significantly affected by climate changes occurring in high-latitudes, with teleconnections resulting from ocean circulation (e.g. changes in the lateral advection of warmer waters). On orbital and glacial-interglacial time scales, long-term SST of IPWP varied coincidently with atmosphere pCO 2 , global sea level and Antarctica atmosphere temperature change (Bolliet et al., 2011;[START_REF] Lea | Climate impact of late quaternary equatorial Pacific sea surface temperature variations[END_REF][START_REF] Lo | Nonlinear climatic sensitivity to greenhouse gases over past 4 glacial/interglacial cycles[END_REF][START_REF] Tachikawa | CO2 radiative forcing and Intertropical Convergence Zone influences on western Pacific warm pool climate over the past 400ka[END_REF], implying tight teleconnections between low-and high-latitudes.

Interestingly, thermocline temperature changes in the IPWP region (reconstructed from the analysis of sub-surface dwelling planktonic foraminifers) appear to be largely decoupled from surface temperature changes. For example, during the Holocene, the SST varied less than 1°C in most of the IPWP region. In contrast, the thermocline water temperature (TWT) declined by 2 -4 °C from early Holocene (~11 ka) to the present. Such thermocline temperature changes have been observed in the western Pacific (Bolliet et al., 2011;Dang et al., 2012;Hollstein et al., 2018), in the middle of Indonesian Archipelago (Fan et al., 2018) and in eastern tropical Indian Ocean (Xu et al., 2008). These observations suggest that the thermocline shoaled along the course of the Holocene over much of the IPWP. This has important consequence on the heat storage capacity of this area. Not only did the temperature declined since the early Holocene, but the thickness of the surface and sub-surface waters in which heat is stored, declined over the course of the Holocene, reducing even more the total heat storage capacity of the IPWP from the early Holocene to present.

In this thesis work, I provide new high resolution Mg/Ca based temperature records form both surface and subsurface water depth in the south sector of IPWP in the eastern tropical ocean side over the last 270, 000 years. These data provide the record of heat evolution in the upper ocean of IPWP, and as well as served as the basis for different studies in the thesis.

The Indonesian Throughflow

At present

The Indonesian Archipelago and its nearly 14,000 islands are located in the middle of the IPWP and constitutes a major gateway between the Pacific Ocean and the Indian Ocean. The Indonesian Archipelago is the only low-latitude pathway for surface and sub-surface waters (down to about 700 m) flowing from the western Pacific Ocean into the eastern Indian Ocean. This major water flow is called Indonesian Throughflow (ITF) (Figure 5). It is the only lowlatitude return branch of the Great Conveyor Belt (AL [START_REF] Gordon | INTER-OCEAN EXCHANGE OF THERMOCLINE WATER[END_REF]. The ITF transport is an important regulator of the meridional overturning of the Indian and Pacific Oceans, and possibly the global thermohaline circulation [START_REF] Gordon | INTER-OCEAN EXCHANGE OF THERMOCLINE WATER[END_REF]. The ITF exports a large amount of heat and freshwater from the Pacific to the Indian Ocean [START_REF] Gordon | Makassar Strait throughflow, 2004 to 2006[END_REF]Gordon, 2005;Sprintall et al., 2009).

The surface flow of the ITF is mainly driven by the pressure gradient between the western Pacific and eastern Indian Ocean and is, therefore, strongly impacted by seasonal changes in wind intensity and direction (Sprintall et al., 2009). Easterly trade winds on the Pacific Ocean blow surface waters, which accumulate on its western side, forming an area with a relatively higher sea level than the eastern Indian Ocean, thus driving surface water flow from the western Pacific Ocean to the eastern Indian Ocean [START_REF] Wyrtki | Indonesian through flow and the associated pressure gradient[END_REF]. On a yearly basis, the ITF transports about 15 Sv (1 Sv = 1 million cubic meters per second) of surface and subsurface water from the Pacific to the Indian Ocean, but with important seasonal changes due changes in the direction and intensity of monsoon-related winds (Sprintall & Révelard, 2014). Although the ITF is regarded as an important component of the Indo-Pacific and global overturning circulation, its transport and variability have been estimated using wind-driven circulation theory. There is still a lack of general understanding about how the deep overturning circulation contributes to the ITF transport.

The primary pathway allowing the Pacific water to enter into the Indonesian seas is the Makassar Strait. As the Makassar Strait is only 680 m deep, only the upper layer water is allowed to pass (Gordon, 2005). Modern observations reveal that ~ 12 Sv of ITF transport takes place through the Makassar Strait, accounting for ~ 80% of the total ITF transport into the Indian Ocean (Gordon, 2005). The ITF enters into the Indian Ocean through three main outflow passages along the Nusa Tengarra archipelago. They are the Lombok Strait (accounting for an annual average of 2.6 Sv), the Ombai Strait (4.9 Sv) and the Timor Passage (7.5 Sv) (Figure 1.5) (Sprintall et al., 2009). After entering the Indian Ocean, a sub-branch of ITF waters travels towards Africa as part of the Indian South Equatorial Current (SEC) and eventually exits the Indian Ocean into the Atlantic Ocean through Agulhas leakage around the southern tip of Africa (Gordon, 2005;Talley, 2013). Another branch of the ITF joins the Leeuwin Current (LC), which travels southward along the west coast of Australia, and reaches as far south as Tasmania (Gordon, 2005). The ITF water is shown to warm the Agulhas Current and the region off the south coast of Java and Sumatra [START_REF] Godfrey | The effect of the Indonesian throughflow on ocean circulation and heat exchange with the atmosphere: A review[END_REF]Gordon, 2005;[START_REF] Oppo | The Great Indo-Pacific Communicator[END_REF]. Thus, ITF is important contributor of a circulation system that transports oceanic heat from the Pacific Ocean to the Indian Ocean and from low latitudes to the high, southern latitudes of Indian Ocean and then the Atlantic Ocean, through the Agulhas leakage. [START_REF] Gordon | Makassar Strait throughflow, 2004 to 2006[END_REF]Sprintall et al., 2009).

In the past

The variability and strength of the ITF transport during the past is still poorly constrained, as there are no direct ways to estimate the past changes in ITF flux. The few existing studies on ITF variability use the thermocline water temperature variations as an indicator. For example, at the inflow region of Makassar Strait, (Fan et al., 2018) found increased thermocline temperature gradient between the north and south of Makassar Strait during the interval 27-24 ka and 19-13.4 ka, these periods were considered as weak ITF, based on the assumption that ITF water from the upstream (that is north of Makassar strait) would warm the downstream (that is south of Makassar strait). Similarly, in the ITF outflow region of Timor sea, Holbourn et al. (2011) reconstructed the thermal gradient between east and west (near to the Indian Ocean)

Timor sea, at thermocline depth by using P. obliquiloculata Mg/Ca derived thermocline water temperature (TWT). Their assumption of ITF effect on the local TWT change is in contrast with the setting in the inflow region of Makassar Strait by Fan et al. (2018), because ITF thermocline water is colder than the Indian Ocean at the same depth, thus ITF at present is cooling the intermediate water mass of the Indian Ocean rather than warm it [START_REF] Gordon | Cool Indonesian throughflow as a consequence of restricted surface layer flow[END_REF]. Thus, ITF from the east side of Timor sea is supposed to cool the thermocline temperature at the west side (near to the Indian Ocean) of Timor sea (Holbourn et al., 2011).

Based on this, Holbourn et al (2011) identified two intervals of 112-86 ka and 56-20 ka as the weak ITF thermocline transport. However, the problem with such an approach, is that both seasonal and inter-annual temperature patterns associated to ITF are partially masked by the superimposition of local changes in temperature related to the development of seasonal upwelling cells forced by monsoon-winds.

As shown above, The ITF is the main low-latitude return branch of the global thermohaline circulation. The heat carried by the ITF from the tropical western Pacific to the eastern tropical Indian Ocean strongly affects local and global climate changes. The magnitude and variability of its transport at different timescales are key factors in the inter-basin transfer of heat and salt and climate teleconnections between low and high latitudes. However, the past change in the strength of ITF were poorly constrained. In this thesis, I provide new evidence for constraining the variability of ITF in terms of both its flux and heat at millennial to orbital timescales.

Intertropical Convergence Zone and Australian-Asian Monsoon

At present

Intertropical convergence zone (ITCZ) is a narrow tropical belt of deep convective clouds centered on average around 6°N north of the Equator. As rainfall is most intense in ITCZ, its position can be identified as the maximum in time-mean precipitation (Schneider et al., 2014).

Seasonally, the ITCZ migrates towards the warmer hemispheres (i.e. the hemisphere that experiences the warm seasons, spring and summer; Schneider et al., 2014). As shown in Figure 1.6, in the IPWP region, ITCZ seasonally migrate from ~10°S during austral summer (i.e. boreal winter) to ~ 20°N during boreal summer. Along with the seasonal migration of ITCZ, local prevailing winds shift in direction (Figure 1.7). The ITCZ movements are what ultimately connect regions belonging to monsoon systems (Global Monsoon). In our studied area, the ITCZ movements are intimately connected with seasonal changes in wind direction and precipitation intensity. Here, Southeast Asia and northern Australia are combined into one monsoonal system, the Malaysian-Australian Monsoon (or Australian-Indonesian monsoon), that differs from others because of the peculiar and somewhat symmetrical distribution of landmasses on both sides of the equator. In this respect, this monsoon is unique. The substantial masses of water between Asia and Australia have a moderating effect on tropospheric temperatures, weakening the summer monsoon. The many islands (e.g., Philippines and Indonesia) provide an infinite variety of topographic effects and play an important role in the spatial variability of precipitations (Figure 1.6). Typhoons that develop within the monsoonal air bring also additional complications. In the study region (between the north Australia and the south Indonesia), during austral winter/boreal summer, dry, cold air blows from Australia to the equator, and southeast winds are prevailing. The ITCZ is shifted to the North of the Equator.

In contrast, during austral summer/boreal winter, wet, warm air blows from the northwest, bringing monsoon rains to the Timor Sea (Figure 1.7). The ITCZ shifts to the South of the Equator, reaching the northernmost par of Australia.

In the past

The strength, range and position of ITCZ dramatically affect the distribution of rainfall in the tropics. The mean position of ITCZ has been thought to migrate latitudinal in the past (Haug et al., 2001;Schneider et al., 2014). Evidence comes mostly from precipitation-sensitive paleoclimate records: land cave stalagmites and ocean sediment cores (e.g. Ayliffe et al., 2013;Carolin et al., 2013;Cheng et al., 2016;Gibbons et al., 2014;Levi et al., 2007;Schneider et al., 2014). Consistent inference from these studies is that the mean position of ITCZ shift southward during the cold Heinrich events during the last glacial period (Schneider et al., 2014). In the IPWP region, a southward migration of ITCZ generally decrease the monsoon rainfall in the north part of equator and increase the local summer monsoon rainfall in the south part of equator (Denniston et al., 2017). From the early Holocene to present, ITCZ exhibited southward migration (Haug et al., 2001;Mohtadi et al., 2011). Along with the shifting of ITCZ through the Holocene, Australian-Indonesian winter monsoon decreased and Australian-Indonesian summer monsoon increased (Ayliffe et al., 2013;Mohtadi et al., 2011).

Consequently, southeast winds were depressed and northwest wind were enhanced in the Australian-Indonesian region. The change in the prevailing winds affected the upwelling intensity and associated primary productivity along the coast Java-Sumatra (Lueckge et al., 2009;Wang et al., 2018). The meridional shift of ITCZ was also thought to vary coherently with Southern Ocean Westerlies and the subtropical front, through which link to the strength of Agulhas leakage and wind-driven upwelling in the Southern Ocean [START_REF] Anderson | Wind-Driven Upwelling in the Southern Ocean and the Deglacial Rise in Atmospheric CO2[END_REF][START_REF] Bard | Migration of the subtropical front as a modulator of glacial climate[END_REF]Beal et al., 2011).

Transient modeling of hydrographic variations across the last deglaciation suggests that past changes in surface ocean d 18 O in the IPWP region were strongly controlled by the average position and dynamics of the ITCZ, reflecting the evolution of the mean state of interhemispheric temperature gradients (Gibbons et al., 2014). This means that long d 18 O record could be used to explore the dynamics of the ITCZ at orbital time scales. So far, long records of such ocean δ 18 O is sparse in IPWP region. In this thesis work, I provide a high-resolution record of surface ocean d 18 O in the IPWP region which is used to study past hydroclimate changes in the IPWP region and explore the mechanisms that control them at different timescales.

Chapter 2 : Material and Methods

This chapter focuses on the material and methods. It summarizes and completes the pieces of information available in the specific chapters built as stand-alone manuscripts for submission into peer-review journals). It provides the readers with additional items about cores MD98-2165 and MD98-2166 and digs more deeply into some methodological aspects that will be superficially discussed in the chapters. In this chapter, I will put also the emphasis on the reconstruction of past water properties (i.e., SST, SSS and TWT).

Material

Core MD98-2165

Sediment core MD98-2165 (9.65S, 118.34E) was retrieved at water depth of 2100 m at the east of Sumba, eastern tropical Indian Ocean, during the Images IV cruise on board French R/V Marion Dufresne in 1998 (Figure 2.1). The core was cut into 1.5m-long sections and each of these sections were split open on board, providing a working half-section and an archive halfsection.

The core is 42.3 m long. There are three sediment gaps between 650 and 677 cm, 1898 and 1909 cm, 3706 and 3723 cm due to coring perturbation. When constructing the age model and converting depths to ages, these gaps are removed from the sub-bottom depth-scale. However, such a "corrected" sub-bottom depth scale can evolve through time when studies reveal the occurrence of disturbed intervals (e.g. turbidities, reworked sediment) that had not been previously noticed. Thus, in order to keep a robust and invariant identification for the samples, each discrete sample has been always labeled with its section number and the depth interval in the section (with reference to the core liner, without applying any correction for gap or sediment disturbance).

The sediment in core MD98-2165 consists of clay-carbonate mud with biogenic and detrital silica. Pyrite is present all over the core. The sediment composition does not vary significantly along the core. The core being located at 2100 m, it is above the modern lysocline depth, which is located at 3800 m in the equatorial Indian Ocean [START_REF] Peterson | Carbonate dissolution in Recent sediments of the eastern equatorial Indian Ocean: Preservation patterns and carbonate loss above the lysocline[END_REF]. The in situ calcitesaturation state of bottom waters (∆[CO 3

2-]), the difference between in situ [CO 3 2-] and saturation state of [CO 3 2-], was calculated by using the CO2Sys software (Lewis et al., 1998) fueled with total alkalinity, temperature, salinity, pH and nutrient values obtained from the nearest GLODAP sites (Sabine, et al., 2005). Estimated ∆[CO 3 2-] is +10 µmol/kg, indicating that bottom water are over-saturated with respect to calcite. Correspondingly, foraminiferal tests are in good preservation all along the core.

Previous analyses performed on this core include 20 accelerator mass spectrometry (AMS) 14 C dates, at several depth intervals between 0 and 800cm (Waelbroeck et al., 2006). These studies also provided high resolution oxygen isotope records from both planktonic and benthic foraminiferal along the whole core (Levi, 2003;Levi et al., 2007;Waelbroeck et al., 2006) and a record of G. ruber Mg/Ca covering the 60 -800cm interval (Levi et al., 2007). These data were used and completed in the present thesis. AMS 14 C dates (Table 2.1), already published by Waelbroeck et al. (2006), suggest that the top 16 cm may have been reworked. In addition, the Mg/Ca data obtained by Levi et al. (2007) revealed anomalously high SST in the top section, down to 70 cm, although the 14 C age at 55 cm appears to coherent with the lower part of the depth-age profile.

In the present thesis, based on AMS 14 C and Mg/Ca arguments, we decided to exclude the upper 110 cm and only study from 110 cm downward. New samples were taken every 4 to 12 cm from the interval 110 -3450 cm. Three planktonic species were hand-picked under a binocular (see below) and the composition in minor and trace elements relative to calcium were measured by Inductively Coupled Plasma-Mass Spectrometry (ICP-MS).

Core MD98-2166

The sediment core MD98-2166 (11.96S, 120.46E) was retrieved at water depth of 3875 m in the northwestern part of the Timor sea, in eastern tropical Indian Ocean, during the Images IV cruise on board French R/V Marion Dufresne in 1998. The core is 34.7 m long. The sediment consists of clays, silts and sands, and is rich in biogenic material (Moreno et al., 2008). Core MD98-2166 is located only about 250 nautical miles (~ 460 km) from core MD98-2165, but is located at a deeper water depth, below the modern lysocline depth (3800 m) in the eastern equatorial Indian Ocean [START_REF] Peterson | Carbonate dissolution in Recent sediments of the eastern equatorial Indian Ocean: Preservation patterns and carbonate loss above the lysocline[END_REF]. The bottom water ∆[CO 3 2-] was estimated by using CO2Sys software (Lewis et al., 1998) and total alkalinity, temperature, salinity, pH and nutrient values obtained from the nearest GLODAP sites (Sabine, et al., 2005). This calculation indicates that bottom water ∆[CO 3 2-] is under-saturated with respect to calcite, at about -11 µmol/kg. This under-saturation likely persisted for most of period covered by the 34.7 m of the sediment core, as shown by the poor preservation of foraminiferal tests and the large number of damaged and fragmented planktonic foraminiferal tests. Thus, in the present thesis, this core is chiefly used for studying the preservation state effect on trace elements in foraminifer shells, in comparison with the shallower, better preserved Core MD98-2165.

In a previous study, Moreno et al. (2008) have obtained a planktonic oxygen isotope record for the whole core and constructed an initial age model. In the present thesis, only the upper 5.4 m of the core were studied. New samples were taken every 10 cm. Three planktonic species were hand-picked under a binocular (see below) and the composition in minor and trace elements relative to calcium were measured by Inductively Coupled Plasma-Mass Spectrometry (ICP-MS).

Methods

Revision of age models

Age model of core MD98-2165

An initial age model for the upper section (55 to 800 cm) was constructed based on accelerated mass spectrometry (AMS) 14 C dates measured from mono-specific G. ruber white (Waelbroeck et al., 2006), and calendar ages were computed by using the 1998 marine calibration curve with a surface reservoir age of 400 years (Waelbroeck et al., 2006). Here, the latest MARINE13 age calibration curve (Reimer et al., 2013) (Paillard et al., 1996). For this tuning exercise, 16 tie points (Table 2.3, Figure 2.3) were inserted and ages were linearly interpolated between these tie points. The average sedimentation rate of MD98-2166 is ~10 cm/ka, ranged between 5 and 25 cm/ka. (Levi, 2003;Levi et al., 2007) and MD98-2166 (Moreno et al., 2008), red vertical bar indicates tie points in Table 2.3. (b) Depth-VS-Age plot by using tie point in Table 2.3, and the variation of sedimentation rates (red line).

Foraminiferal geochemistry proxies used in this work

Foraminifera are single-cell marine protozoans with calcareous shells (also called tests). They are divided into two types: benthic foraminifera, which live in the bottom of oceans, and planktonic foraminifera, which live in the surface and subsurface layers of the oceans.

Foraminifers construct their tests either by using dissolved carbonate and calcium ions from ambient seawater (planktonic and benthic foraminifers) or by agglutinating solid particles from the sediment (some benthic foraminifers). Species that precipitate their tests from dissolved Ca 2+ and CO 3 = and precipitate calcite or, more rarely, aragonite. The elemental and isotopic composition of foraminiferal carbonate tests depends upon the physical and chemical properties of ambient seawater in which the organisms develop. Changes of these physical and chemical properties carry environmental and climatic signals. In the present work, two types of proxies are used: the stable isotopic composition of foraminiferal tests and their trace elements composition.

Stable oxygen isotopes in seawater and foraminiferal tests

Stable oxygen isotopic composition of foraminiferal tests, expressed as δ 18 O cal (Equation 2.1)

with respect to an international reference standard, is certainly the most widely used among paleoceanography proxies. As ocean flows, in addition to the evaporation/precipitation, at a certain ocean site, the seawater δ 18 O sw will be largely affected by the δ 18 O sw of the source water. During the past, on millennial to glacial-interglacial timescales, the melting and growth of ice sheets dramatically influenced the average seawater δ 18 O sw , because Greenland and Antarctic ice sheets are large reservoir of water with extremely low δ 18 O values, ranging from -30‰ to -50‰ (EPICA Community Members) with respect to the average, reference value of 0‰ for modern ocean water (LeGrande & Schmidt, 2011). Thus, future melting of ice sheets will decrease the seawater δ 18 O sw . Over the last deglaciation, from the Last Glacial Maximum (~20ka) to the early Holocene (~11ka), the melting of ice sheets leads to ~120 m sea-level rise which was accompanied by a ~1.2‰ decrease of global mean seawater δ 18 O sw (Adkins et al., 2002;Waelbroeck et al., 2002). Thus, over the past 2.4 Ma of the "Ice Age" world [START_REF] Imbrie | The Orbital Theory of Pleistocene Climate: Support frim a Revised Chronology of the Marine d18O Record[END_REF], waxing and waning of large, high latitude ice sheets, paced by orbital changes in insolation, lead to quasi-periodic changes in the average δ 18 O composition of the global ocean [START_REF] Nicholas | The last interglacial in the marine and terrestrial records[END_REF]. These changes, recorded in fossil planktonic shells, provide the backbone for the stable isotopic stratigraphy, a powerful stratigraphic tool [START_REF] Bassinot | The Astronomical Theory of Climate and the Age of the Brunhes-Matuyama Magnetic Reversal[END_REF][START_REF] Imbrie | The Orbital Theory of Pleistocene Climate: Support frim a Revised Chronology of the Marine d18O Record[END_REF]LE Lisiecki & Raymo, 2005;[START_REF] Shackleton | The 100,000-year ice-age cycle identified and found to lag temperature, carbon dioxide, and orbital eccentricity[END_REF]; N. J. [START_REF] Shackleton | Oxygen isotope and palaeomagnetic stratigraphy of Equatorial Pacific core V28-238: Oxygen isotope temperatures and ice volumes on a 105 year and 106 year scale[END_REF].

But seawater composition is not the only parameter that drives the δ 18 O cal. composition of foraminiferal calcite. The seawater δ 18 O sw acts together with seawater temperature (Equation 2.2). Indeed, during the precipitation of calcite, thermodynamic effects result in a temperaturerelated isotopic fractionation (calcite δ 18 O becoming lighter with increasing precipitation temperature). First evidenced by [START_REF] Urey | The thermodynamic properties of isotopic substances[END_REF] for inorganic calcite, this temperature fractionation effect was later confirmed for marine bio-calcifying organisms, making it possible the development of stable isotopic paleo-thermometry [START_REF] Emiliani | Pleistocene temperatures[END_REF]. An example of temperature, δ 18 O sw and δ 18 O cal relationship is displayed in Equation 2.2, which was established by Bemis et al. (1998) by culture experiments on the planktonic species Orbulina universa.

δ ,-$ %*&L&7*: MNO$P = 0.27 + S(°C) -16.5 4.8 + δ ,-$ \&)\]7* M^_` (Equation 2.2)

The paleo-thermometer approach is made more complex by the fact that non-thermodynamic effects (so-called "vital effects") can modify the temperature-related δ 18 O fractionation [START_REF] Duplessy | Differential Isotopic Fractionation in Benthic Foraminifera and Paleotemperatures Reassessed[END_REF][START_REF] Ravelo | Oxygen Isotopic Composition of Multiple Species of Planktonic Foraminifera: Recorders of the Modern Photic Zone Temperature Gradient[END_REF][START_REF] Spero | Experimental determination of stable isotope variability in Globigerina bulloides: implications for paleoceanographic reconstructions[END_REF]. These vital effects are species-dependent, potentially reflecting kinetic effects and non-equilibrium reactions mediated by biological mechanisms. In surface planktonic foraminifers, for instance, vital effects can be related to the photosynthetic activity of algal symbionts. Species containing algal symbionts were found to have higher calcification rates, which may result in 0.35 -0.5‰ depletion in calcite δ 18 O cal relative to species without symbionts [START_REF] Ravelo | Oxygen Isotopic Composition of Multiple Species of Planktonic Foraminifera: Recorders of the Modern Photic Zone Temperature Gradient[END_REF][START_REF] Spero | Experimental determination of stable isotope variability in Globigerina bulloides: implications for paleoceanographic reconstructions[END_REF].

Mg/Ca and other Elements/Ca in foraminiferal tests

Various elements can be incorporated into foraminiferal tests either in substitution of calcium, or as impurities, within the defaults of the crystal lattice. The proportions of these elements, expressed as ratios relative to calcium (element/Ca) are related to the physiochemical properties of seawater in which foraminifera make their tests (e.g., [START_REF] Allen | Trace element proxies for surface ocean conditions: A synthesis of culture calibrations with planktic foraminifera[END_REF][START_REF] De Nooijer | Biomineralization in perforate foraminifera[END_REF]. One of the most important element/Ca ratios used in paleoceanographic studies and which will be the main tool of this thesis is Mg/Ca.

Laboratory culture experiments (e.g., [START_REF] Allen | Trace element proxies for surface ocean conditions: A synthesis of culture calibrations with planktic foraminifera[END_REF], sediment trap experiments (e.g, Anand et al., 2003;[START_REF] Gray | Nonthermal Influences on Mg/Ca in Planktonic Foraminifera: A Review of Culture Studies and Application to the Last Glacial Maximum[END_REF]Mohtadi et al., 2009) and surface sediments studies (e.g., [START_REF] Lea | Climate impact of late quaternary equatorial Pacific sea surface temperature variations[END_REF]Mohtadi et al., 2011) all reveal that planktonic foraminiferal Mg/Ca is exponentially correlated to the ambient seawater temperature, as expressed in Equation 2.3.

Oa bc = ` * d e * f (Equation 2.3)

In this equation, T is temperature, B and A are pre-exponential and exponential constants which can be determined by studying modern foraminiferal samples (from core tops or cultured)

together with instrumental record of temperature data (e.g. Anand et al., 2003). Although some authors have suggested to use multi-species, empirical relationship, several studies have highlighted small differences between Mg/Ca-T° relationship, clearly suggesting the possibility that vital effects also play a role (Anand et al., 2003). In addition to temperature, non-thermal factors such as salinity and pH have been reported to exert secondary effects on planktonic foraminiferal Mg/Ca composition (e.g. [START_REF] Kisakurek | Controls on shell Mg/Ca and Sr/Ca in cultured planktonic foraminiferan, Globigerinoides ruber (white)[END_REF][START_REF] Hönisch | The influence of salinity on Mg/Ca in planktic foraminifers -Evidence from cultures, core-top sediments and complementary δ18O[END_REF]Mathien-Blard and Bassinot, 2009;[START_REF] Gray | Nonthermal Influences on Mg/Ca in Planktonic Foraminifera: A Review of Culture Studies and Application to the Last Glacial Maximum[END_REF]. By compiling published laboratory culture experiments data obtained on four planktonic species (Globigerinoides ruber, Trilobatus sacculifer, Globigerina bulloides, and Orbulina universa), [START_REF] Gray | Nonthermal Influences on Mg/Ca in Planktonic Foraminifera: A Review of Culture Studies and Application to the Last Glacial Maximum[END_REF] concluded that Mg/Ca-sensitivity to salinity is 4% per salinity unit (PSU), and to pH is -5 to -9% per 0.1 pH unit. That Mg/Ca-sensitivity to salinity was obtained in large salinity range of 30 -45 PSU, but below a threshold of 35 PSU, salinity change has no significant effect on G. ruber Mg/Ca (e.g., [START_REF] Hönisch | The influence of salinity on Mg/Ca in planktic foraminifers -Evidence from cultures, core-top sediments and complementary δ18O[END_REF][START_REF] Gray | The effects of temperature, salinity, and the carbonate system on Mg/Ca in Globigerinoides ruber (white): A global sediment trap calibration[END_REF][START_REF] Gray | Nonthermal Influences on Mg/Ca in Planktonic Foraminifera: A Review of Culture Studies and Application to the Last Glacial Maximum[END_REF].

When studying fossil foraminiferal tests, two post-depositional processes -dissolution and diagenetic overgrowth -are able to bias Mg/Ca temperatures. Mg/Ca ratios are distributed nonhomogeneously in foraminiferal tests (e.g. Sadekov et al., 2009;[START_REF] Evans | LA-ICPMS Ba/Ca analyses of planktic foraminifera from the Bay of Bengal: Implications for late Pleistocene orbital control on monsoon freshwater flux[END_REF]. It has been shown that dissolution tends to remove preferentially the Mg-rich parts of the test due to the poorer crystallinity of the calcite lattice (Nouet & Bassinot, 2007), and thus in situ dissolution on seafloor generally decreases the Mg/Ca ratio of foraminifers (e.g., Brown & Elderfield, 1996;Regenberg et al., 2014).

Fossil foraminiferal shells always contain contaminants which may be rich in Mg. Common

Mg contaminants are clay minerals, organic matters and ferromanganese oxides (Barker et al., 2003;Pena et al., 2008). In order to reconstruct accurate Mg/Ca temperatures, it is crucial to clean foraminiferal tests to remove Mg-rich contaminants. There exists, basically, two main approaches for this cleaning: one including a reductive step to remove oxide coatings (Boyle & Keigwin, 1985), and one approach without that reductive step (e.g. Barker et al., 2003).

Including or not this reductive cleaning step will result in different Mg/Ca values (e.g., [START_REF] Rosenthal | Accurate estimation of sea surface temperatures using dissolution-corrected calibrations for Mg/Ca paleothermometry[END_REF]Barker et al., 2003). Generally, including a reductive step will lead to a decrease in Mg/Ca ratios (Barker et al., 2003;Yu et al., 2007;Bian & Martin, 2010).

Yet, these literatures are rather confusing regarding this cleaning effect, some authors claiming that there are substantial differences in Mg/Ca due to the cleaning, others claiming that the effects are null or can be neglected. In addition, the amplitude of the cleaning effect is controversial for a given species. In this PhD work, I devoted some time to better understand the effects and potential bias of cleaning procedures as insuring coherency between data provided by different laboratories that do not use the same methodological approaches is crucial when trying to build comprehensive datasets of past ocean changes. The chapter 3 will present these results about cleaning bias and possibilities to correct for it. Other elemental ratios such as B/Ca, Li/Ca, Sr/Ca, Ba/Ca have been used to derive pieces of information about various physic-chemical parameters in the ocean [START_REF] Allen | Trace element proxies for surface ocean conditions: A synthesis of culture calibrations with planktic foraminifera[END_REF][START_REF] Delaney | Lithium in foraminiferal shells: implications for hightemperature hydrothermal circulation fluxes and oceanic crustal generation rates[END_REF][START_REF] Marr | Ecological and temperature controls on Mg/Ca ratios of Globigerina bulloides from the southwest Pacific Ocean[END_REF][START_REF] Weldeab | 155,000 years of West African monsoon and ocean thermal evolution[END_REF]. For example, B/Ca in planktonic foraminifera was used as a proxy for seawater pH [START_REF] Yu | B/Ca in planktonic foraminifera as a proxy for surface seawater pH: B/Ca IN PLANKTIC FORAM AS A PROXY FOR pH[END_REF] and in benthic foraminifera was considered as a proxy of deep water carbonate saturation state change (Yu & Elderfield, 2007), Ba/Ca in planktonic foraminifera was used as a proxy of river runoff [START_REF] Weldeab | 155,000 years of West African monsoon and ocean thermal evolution[END_REF], Li/Ca was thought to reflect the seawater carbonate ion concentration [START_REF] Hall | Li/Ca in multiple species of benthic and planktonic foraminifera: thermocline, latitudinal, and glacial-interglacial variation[END_REF]. In this thesis, these elements (Li, B, Sr, Ba, U, Mn, Fe and Al) were

simultaneously measured with Mg and Ca from three planktonic foraminiferal species -G.ruber, P.obliquiloculata and G. menardii, in the two sediment cores -MD98-2165 and MD98-2166. Those elements were not used in the present work except for Mn, Al and Fe (used to decipher the efficiency of cleaning methods in removing detrital or oxide contaminants). All

Elements/Ca ratios results are provided in large figures in the Annex.

Foraminiferal species used in this work

As seen above, analyzing element ratios and isotopic composition from fossil foraminiferal tests picked from marine cores, make it possible to reconstruct past marine conditions through time. One of the advantages of using biogeochemistry proxy from foraminiferal shell is the ability to reveal marine conditions changes from different water layers. In this thesis, three planktonic foraminifera species with different dwelling depths were studied.

Globigerinoides ruber (white)

G. ruber is a mixing layer species. G. ruber has two main morphotypes, one is called G. ruber sensu stricto (s.s.), another is called G. ruber sensu lato (s.l.) (Figure2.7; [START_REF] Wang | Isotopic signals in two morphotypes of Globigerinoides ruber (white) from the South China Sea: implications for monsoon climate change during the last glacial cycle[END_REF]. In the south China sea, shell geochemistry (e.g., δ 18 O) in the two morphotypes are different, as G. ruber (s.l.) has a deeper habitats depth than G. ruber (s.s.) [START_REF] Wang | Isotopic signals in two morphotypes of Globigerinoides ruber (white) from the South China Sea: implications for monsoon climate change during the last glacial cycle[END_REF]. However, sediment trap samples and surface sediment studies along the coast of south Indonesian (study region of this thesis) suggest the same habitat depth for both G. ruber (s.s.) and G. ruber (s.l.) of 0 -30m, and no significant difference in shell geochemistry (e.g., δ 18 O and Mg/Ca) (Mohtadi et al., 2009(Mohtadi et al., , 2011)). The consistent habitat depth and hence same geochemistry character between G. ruber (s.s.) and G. ruber (s.l.) were also reported in other different ocean sites by studying sediment trap samples (e.g., [START_REF] Thirumalai | Globigerinoides ruber morphotypes in the Gulf of Mexico: A test of null hypothesis[END_REF][START_REF] Gray | The effects of temperature, salinity, and the carbonate system on Mg/Ca in Globigerinoides ruber (white): A global sediment trap calibration[END_REF]. Therefore, in this thesis, G. ruber (s.s.) and G. ruber (s.l.) were mixed together for the convenient of obtaining enough tests. But only the morphotype of number 4-6 as shown in Figure 2.7 were selected as G. ruber (s.l.), the morphotype of number 7-9 in Figure 2.7 was not included. In the study region of this thesis, sediment trap experiments show that although enhanced G. ruber flux rates exist during the crushed between two glass plates under a binocular and thoroughly mixed to insure optimum homogeneity of samples aliquots. Samples were divided as evenly as possible in two subsamples (Group A and Group B) and transferred to acid-cleaned 500 µl centrifuge tubes for further cleaning.

For the rest of samples which need to be cleaned by only one same method (MD98-2165: 745 -3450 cm), 30 individuals of G. ruber and P. obliquiloculata in the size fraction 250 -355 µm were hand-picked. Samples were then gently crushed between two glass plates under a binocular and transferred to acid-cleaned 500 µl centrifuge tubes for further cleaning. There samples were attributed to Group A.

Cleaning

Group A and Group B sub-samples were cleaned following the protocol presented in Barker et al. (2003), and a reductive step was added for samples from Group B (Figure 2.10). Here is a brief description of cleaning, for further details, please refer to Barker et al. (2003). Samples were ultrasonically cleaned in Milli-Q water (5 times) then in methanol (2 times) for removing clay minerals. Two additional rinses in Milli-Q water were performed for removing any remaining methanol. Samples were then checked under a binocular and any coarse silicates and dark-colored particles were removed before continuing with the oxidative cleaning step. For Group B samples, a reductive step was added prior to the oxidative cleaning. For this reductive step, I followed the protocol initially developed by Boyle and Keigwin (1985) and later used or slightly modified in other studies (eg. Martin & Lea, 2002;Yu et al., 2007;Brian & Martin, 2010). Samples were put in 0.12M citrate buffered with ammonium hydroxide and hydrazine solution in a boiling water bath for 30 minutes. Samples were ultra-sonicated every 5 minutes for promoting the reaction. Samples were then rinsed twice with Milli-Q water. For the rest of cleaning procedure, samples Group A and Group B were treated the same way. They were soaked into alkali-buffered 1% H 2 O 2 solution in a boiling water bath for 15 minutes (1 time).

Every 3 minutes, samples were put into an ultrasonic bath for a few seconds for releasing CO 2 bubbles building up during the oxidative reaction. Samples were then rinsed twice with Milli-Q water for removing remaining oxidative reagent. The cleaning protocol ends up with a rapid leaching in a diluted nitric acid solution (HNO 3 0.001M) for removing any adsorbed contaminants, followed by thorough rinsing with Milli-Q water. All the water was removed through pipetting and samples left in the centrifuge tubes. After cleaning, each sample was dissolved with 0.075M HNO 3 and then centrifuged for removing undissolved particulate

Standard solutions for Elements/Ca calibration

The elements output from the ICP-MS are with the unit of count per second (c/s). The Elements/Ca ratios were then converted to mmol/mol or µmol/mol by using calibration equations constructed by simultaneously measured standard solutions series in which the elements concentration are known (Figure 2.12). The standard solutions were made gravimetrically starting from commercial single element standard solutions in an ultra-clean laboratory in Geops, this standard solution series is called 'Mother' standard solution. This mother standard solution series from Std-J1 to Std-J6 are characterized by super high similar Ca concentration with increased concertation in other elements. The Std-J0 is used as the 'Blank'

for monitoring other elements as the contaminates from the matrix Ca. A theoretic value is shown for each other elements from the Ca matrix. But in practice, the Std-J0 as a blank is subtracted directly from Std-J1 to Std-J6 when build the calibration equations between the machine output (c/s) and the concentration ratios (e.g., mmol/mol). The similar Ca concentration can avoid the potential matrix effect, although such effect is significantly smaller in ICP-MS than the ICP-AES (Yu et al., 2005). The super high Elements/Ca concentration in the mother solution can reduce the risk of contamination and thus the mother standard solutions can be stored and used for many years which is important to keep the coherence and consistency of the measurements. For daily working, a so called daily working standard solution series are used, which are diluted from the mother standard solutions to make the Ca concentration to match with the foraminiferal sample solutions. Examples for the daily working standard solution are given in Figure 2.12. Mg/Ca, Sr/Ca, B/Ca, Li/Ca, Cd/Ca, Ba/Ca, and U/Ca are determined by using daily standard solution (series 1). Mn/Ca, Fe/Ca, Al/Ca, and Ca concertation are determined by daily standard solution (series 2). Cleaning method impact on the Mg/Ca of three planktonic foraminifera species: a downcore study along a depth transect Xiaolei PANG 1,2* , Franck BASSINOT 1 , Sophie SEPULCRE 2 1 Laboratoire des Sciences Du Climat et de l'Environnement, LSCE/IPSL, CEA-CNRS-UVSQ, Université Paris-Saclay, Gif-sur-Yvette, France 2 GEOPS, Université Paris-Sud, CNRS, Université Paris-Saclay, Rue du Belvédère, 91405, Orsay, France.

In preparation

Key Points:

• We performed a downcore comparison of Mg/Ca ratios measured on three planktonic foraminifer species cleaned with reductive and non-reductive cleaning methods.

• Partial dissolution induced by the reductive cleaning step results in a Mg/Ca decrease of ~6% in average compared to the non-reductive cleaning, equivalent to a ~ 0.7°C decrease in reconstructed temperatures.

• Although one can correct for different cleaning methods using this 6% correction factor, our result show that cleaning-induced Mg/Ca differences are species-dependent and are modulated by preservation state of the tests to some extent, suggesting that specific correction factors should be preferred when possible.

• non-reductive cleaning approach should be preferred for building inter-laboratory homogeneous datasets of Mg/Ca-derived temperatures, and the reductive step should be only included when severe oxide coating is observed.

Introduction

Foraminiferal Mg/Ca ratios are used for reconstructing past seawater temperatures (e.g., Anand et al., 2003;;Xu et al., 2010;Elderfield et al., 2012;Fan et al., 2018). In order to reconstruct accurate Mg/Ca temperatures, it is crucial to clean foraminiferal tests to remove Mg-rich contaminants. Common Mg contaminants are clay minerals, organic matter and ferromanganese oxides (Barker et al., 2003;Pena et al., 2008). The removal of clay minerals and organic matter are the most important steps in Mg/Ca cleaning (Barker et al., 2003) and are always included in cleaning methods. In order to remove oxides, a reductive step must be added to the cleaning protocol, with the use of a hydrous hydrazine solution mixed with citrate acid and buffered with ammonia (Boyle & Keigwin, 1985;Martin & Lea, 2002;Yu et al., 2007).

Benefits and pitfalls of removing oxides with such a chemical treatment have been widely debated over these last years, some studies including this reductive step, whereas others preferring not to (Martin & Lea, 2002;Barker et al., 2003;Yu et al., 2007;Bian & Martin, 2010). Indeed, despite the ammonia buffering, several studies have provided evidence of dissolution of foraminiferal tests due to this reductive cleaning, resulting in a decrease of Mg/Ca ratios compared to those obtained from a cleaning protocol without a reductive step (e.g. Rosenthal et al., 2004;Johnstone et al., 2016). According to several authors, this decrease in Mg/Ca reflects the inhomogeneous distribution of Mg in the tests and the preferential dissolution of Mg-enriched parts during the reductive cleaning (Barker et al., 2003;Yu et al., 2007). However, in some sedimentary environments, not removing contaminant oxides can result in erroneously high Mg/Ca values (Pena et al., 2005(Pena et al., , 2008;;Weldeab et al., 2006).

Paleoceanographers are caught, therefore, between a rock and a hard place and no general consensus has yet emerged concerning the cleaning protocol to be used. This lack of consensus can have significant impacts on paleoceanographic reconstructions because, for instance, building databases of past ocean temperatures requires coherent datasets. Addressing interlaboratory biases such as the potential impact of cleaning protocol and how to correct for this are thus key issues to homogenize the published Mg/Ca results.

Several studies have tried to address quantitatively the effects of reductive cleaning on foramininferal Mg/Ca but have reached contrasted conclusions. As part of an inter-laboratory comparison exercise including a multi-species study, Rosenthal et al., (2004) The present study aims at unraveling the species and preservation dependencies of Mg/Caoffsets resulting from using non-reductive and reductive cleaning methods. For this purpose, we picked three planktonic foraminifer species (G. ruber, G. menardii and P. obliquiloculata)

exhibiting different sensitivities to dissolution on two nearby cores retrieved at different water depths in the eastern tropical Indian Ocean.

2.Material and Methods

We selected two adjacent sediment cores retrieved from different water depths in the western Timor Sea (Figure 1): core MD98-2165 (9.65° S, 118.34° E) located at 2100m of water depth, and core MD98-2166 (11.96° S, 120.46° E) located at 3875m of water depth. At the two sites, we calculated the calcite-saturation state of bottom waters (∆[CO 3 2-]) -the difference between in situ [CO 3 2-] and saturation state of [CO 3 2-] -using the CO2Sys software (Lewis et al., 1998) fueled with total alkalinity, temperature, salinity, pH and nutrient values obtained from the nearest GLODAP sites (Sabine, et al., 2005). Estimated ∆[CO 3 2-] is about +10 µmol/kg at the water depth of core MD98-2165 and -11 µmol/kg at the water depth of core MD98-2166. As expected, the difference between the two sites and the under-saturation at site MD98-2166 suggest a much stronger dissolution effect on foraminifer tests at the deeper site. This is in agreement with the higher proportion of foraminifer fragments observed in samples from the deeper core MD98-2166.

62

For core MD98-2165, we focused on the upper part of the core down to 779 cm. Age model of the upper part of MD98-2165 was developed by Waelbroeck et al. (2006) based on 20 14 C dates.

Three anomalous 14 C values (inversion) suggested that the upper 110 cm of the sedimentary are allochtonous and were deposited rapidly (Waelbroeck et al., 2006). We focused our study, therefore, on the interval 110-779cm, covering the time interval from ~ 21 ka Cal BP (Last Glacial Maximum) to the Late Holocene (~2 ka Cal BP). Samples for the present study were taken every 5 to 10 cm (corresponding to time a resolution ranging from 140 to 280 yr).

For core MD98-2166, we studied the upper sedimentary section down to 540 cm. We sampled the sediment series every 10cm. The age model for MD98-2166 was developed by Moreno et al. (2008) We picked three planktonic foraminifera species for this study: G. ruber (white), P.

obliquiloculata and G. menardii. G. ruber and P. obliquiloculata are two species commonly used in the Indo-Pacific Warm Pool region for reconstructing surface and upper thermocline water temperatures, respectively (e.g. Xu et al., 2010;Bolliet et al., 2011;Fan et al., 2018;Wang et al., 2018;). G. menardii is a deep-dwelling species, which has been used to reconstruct past changes in the thermocline structure (Regenberg et al., 2010). These three species have different textures and morphological structures. P. obloquiloculata is characterized by a glassy, outer crust (Steinhardt et al., 2015); G. menardii displays a large and thick surrounding keel and a dense outer crust (Regenberg et al., 2010); G. ruber has a thin and porous test. The outer crusts (P.o and G.m) and the keel of G. menardii are characterized by lower Mg/Ca ratios compared to inner layer parts (Sadekov et al., 2005;Sadekov et al., 2009;Steinhardt et al., 2015), and are known to be more resistant to dissolution occurring on the sea floor or during the cleaning in laboratory (Brown & Elderfield, 1996;Johnstone et al., 2016). Because of these characteristics, P. obliquiloculata and G. menardii are dissolution-resistant species, whereas G.

ruber tests are more easily removed by dissolution in deep marine sediments.

For each sample, about 60-80 individuals of G. ruber and P. obliquiloculata were hand-picked from the 250-355 µm size fraction and about 30 individuals of G. menardii were picked from the 355-450 µm size fraction. The samples were gently crushed between two glass plates under a binocular and thoroughly mixed to insure optimum homogeneity of sample aliquots. Samples were then divided as evenly as possible to obtain two sub-samples (Group A and Group B) and transferred to acid-cleaned, 500 µl centrifuge tubes for cleaning.

Group A and Group B sub-samples were cleaned following the protocol presented in Barker et al. (2003), and a reductive step was added for samples from Group B. Here is a brief description of cleaning, for further details, please refer to Barker et al. (2003). Samples were ultrasonically cleaned in Milli-Q water (5 times) then in methanol (2 times) for removing clay minerals. Two additional rinses in Milli-Q water were performed for removing any remaining methanol.

Samples were then checked under a binocular and any coarse silicates and dark-colored particles were removed before continuing with the oxidative cleaning step. For Group B samples, a reductive step was added prior to the oxidative cleaning. For this reductive step, we followed the protocol initially developed by Boyle and Keigwin (1985) and later used or slightly modified in other studies (eg. Martin & Lea, 2002;Yu et al., 2007;Brian & Martin, 2010). Samples were put in 0.12M citrate buffered with ammonium hydroxide and hydrazine solution in a boiling water bath for 30 minutes. Samples were ultra-sonicated every 5 minutes for promoting the reaction. Samples were then rinsed twice with Milli-Q water. For the rest of cleaning procedure, samples A and B were treated the same way. They were soaked into alkalibuffered 1% H 2 O 2 solution in a boiling water bath for 15 minutes (1 time). Every 3 minutes, samples were put into an ultrasonic bath for a few seconds for releasing CO 2 bubbles building up during the oxidative reaction. Samples were then rinsed twice with Milli-Q water for removing remaining oxidative reagent. The cleaning protocol ends up with a rapid leaching in a diluted nitric acid solution (HNO 3 0.001M) for removing any adsorbed contaminants, followed by thorough rinsing with Milli-Q water. All the water was removed through pipetting and samples left in the centrifuge tubes.

Prior to analysis, each sample was dissolved in 0.075M HNO 3 and then centrifuged for removing any undissolved particulate impurities. The supernatant solution was transferred to a new acid-cleaned, 500 µl centrifuge tube. Mg/Ca ratios were measured by an inductively coupled plasma mass spectrometry (ICP-MS, Analytik Jena PlasmaQuant MS Elite). Aliquots of samples were first analyzed for Ca concentrations and the remaining solutions were diluted to reach a Ca concentration of ~100 ppm in order to minimize possible matrix effects. A standard solution was analyzed every 3 samples to monitor the instrument stability and correct for drifts (Yu et al., 2005). The long-term reproducibility of Mg/Ca measurements estimated from replicate analyses of a standard solution (Mg/Ca=3.69 mmol.mol -1 ) is ±1.4% (n=144 1s relative standard deviation). In each run, the Group A and Group B subsamples of the same initial sample were analyzed sequentially for insuring a better inter-comparison of pairedresults. By doing this, we aimed at minimizing potential differences in Mg/Ca ratios between paired-subsamples that could result from instrumental instabilities. Together with Mg and Ca,

we measured Al and Fe, to check for potential contamination by clay minerals (e.g., Barker et al., 2003), and we measured Mn, to detect the presence of oxide coating.

Results

Contaminants

The low levels of Al (mean value ~ 0.01 mmol.mol -1 , maximum values < 0.2 mmol.mol -1 ) indicate that the cleaning procedure successfully removed any clay phase in the foraminifer samples. No relationships exist between higher concentrations of Al and high Mg/Ca values.

Overall, Mn/Ca ratios remain low, ranging from 0.07 mmol.mol -1 (average value obtained for reductively cleaned P. obliquiloculata in core MD98-2166) up to a maximum value of 0.47 mmol.mol -1 (average value obtained for G. ruber specimens cleaned without a reductive step in core MD98-2165). Mn/Ca values clearly show the impact of reductive cleaning on oxide removal. Adding a reductive step results in a systematic 50 to 70% drop in Mn/Ca values in B sub-samples compared to the A group aliquots cleaned with the non-reductive protocol. For instance, average Mn/Ca in G. ruber from core MD98-2166 drops from 0.47 mmom.mol -1 in the non-reductively-cleaned, A-batch, to 0.2 mmol.mol -1 in the reductively cleaned, B-batch.

Mg/Ca results

G. ruber

For G. ruber, average Mg/Ca values range between ~ 3.2 and 4.7 mmol.mol -1 (Figure 2a) Holocene evolution, suggesting the regional homogeneity and similar temporal evolution of sea-surface temperatures at two nearby sites located at the exit of the Timor Sea.

P. obliquiloculata

The P. obliquiloculata Mg/Ca record from core MD98-2165 displays a clear deglaciation increase, with glacial values around 2 mmol.mol -1 , increasing to ~ 2.7 mmol.mol -1 at the beginning of the Holocene (Figure 3a). Then, Mg/Ca values decrease upward to reach ~ 2.1 mmol.mol remaining very similar to glacial values, around 1.8 mmol.mol -1 . Lower Mg/Ca in Core MD98-2166 may partly result from the poorer preservation at this deeper site. We interpret the different time-evolution of the two P. obliquiloculata Mg/Ca records as resulting from the regional complexity of the thermocline structure and its specific evolution at the two sites. This contrasts with the coherent evolution of sea-surface conditions at the two sites revealed by the similarity of G. ruber Mg/Ca records.

As was already observed for G. ruber, Mg/Ca ratios obtained on aliquots treated with the reductive step are lower in average than Mg/Ca ratios obtained on aliquots cleaned without the reductive step. Those differences show, however, a higher degree of scattering for P.obliquiloculata than for G. ruber; the Mg/Ca differences between the two cleaning methods averaging 4.4 ± 6.8% (1s, n=71) for core MD98-2165, and 6.3 ± 8.8% (1s, n=46) for the deeper core MD98-2166 (Figure 3c-d). As can be observed in Figure 3, part of that variability rests on a few "odd" samples for which reductively cleaned Mg/Ca are higher than the non-reductively cleaned aliquots. The Mg/Ca differences do not display a clear pattern with age (or depth) in cores MD98-2165 and MD98-2166 (Figure 3 (c) and (d)).

G. menardii

For G. menardii, averaged Mg/Ca values range between ~ 1.8 mmol.mol -1 and ~2.7 mmol.mol - 1 in core MD98-2165, and between ~ 1.3 and 2.2 mmol.mol -1 in core MD98-2166 (Figure 4ab). No clear Mg/Ca evolution corresponding to the deglaciation can be seen in the records, potential glacial/interglacial structure being overshadowed by short-scale, high amplitude oscillations and rapid excursions. As was already observed for P. obliquiloculata, G. menardii

Mg/Ca values are significantly lower in core MD98-2166 than in the shallower core MD98-2165 and their temporal evolutions are different too. As was concluded for P. obliquiloculata, the lower Mg/Ca values in G. menardii from Core MD98-2166 may partly result from the poorer preservation at this deeper site. However, we interpret the different time-evolution of the two P. obliquiloculata Mg/Ca records as resulting from the regional complexity of the thermocline structure and its specific evolution at the two sites, which result in different temperature conditions recorded at our two sites.

Similarly, to what was observed for the two previous planktonic species, Mg/Ca ratios obtained with the reductive cleaning protocol (blue curves) are usually lower than Mg/Ca ratios obtained without the reductive cleaning protocol (red curves). The variance is higher than for the two other species, though. In the shallow core MD98-2165, the Mg/Ca cleaning differences average 8.5 ± 6.3% (1s, n=48), whereas they average 3.7% with a particularly high variance of ±11.1% Overall, when considering the three planktonic species together, Mg/Ca ratios from group B (with reductive step) are significantly reduced compared to those from group A (Figures 2,3 and 4) Mg/Ca ratios from group B aliquots being lowered by 6 ± 7% (1h, n=328) compared to aliquots from group A.

Discussion

Statistical robustness of our estimates of cleaning-related Mg loss

Considering the rather high scattering of datasets, it is important to decipher first whether we can confidently compare and rank the Mg/Ca %loss between the three species and between the two sites. Shapiro-Wilk tests indicate that Mg/Ca differences between paired A-B aliquots are random in essence, showing normal distributions. Thus, we are allowed to run t-test to check for the statistical robustness of similarity/dissimilarity between average Mg/Ca losses. Prior to performing those tests, we run F-tests to decipher whether variances of the paired datasets we want to compare are similar or dissimilar, which made it possible for us to select the most appropriate t-test (i.e. Student test for equal variance cases, Welch's t-test for unequal variance cases). Those t-tests show that all the average Mg/Ca% differences obtained for the three species and at the two sites are significantly different (at a 95% confidence level), except those from G. ruber (4.1%) and G. menardii (3.7%) in core MD98-2166, which cannot be significantly discriminated from one another.

Analytical scattering, instrumental stability, inhomogeneity of foraminiferal tests

Despite careful sample preparation and analytical precautions (i.e. sequential ICP-MS analyses of paired A and B aliquots, drift correction), Mg/Ca differences show scattering and variability, which are of the same order of magnitude and sometimes higher than the averaged offsets, and which are much larger than long-term instrumental uncertainty (i.e. ±1.4%, 1s). The lowest variability is observed for G. ruber, which shows a standard deviation of ±3.5% (1s) around the 6.2% mean difference at site MD98-2165 (Figure 2c) and shows a standard deviation of ± 4.6% (1s) around the 4.1% mean at site MD98-2166 (Figure 2d). On the other side of the spectrum, the highest variance is observed for G. menardii, with a standard deviation of ±6.3%

(1s) around a mean value 8.5% at site MD98-2165 (Figure 4c), and a staggering standard deviation of ±11.1% (1s) around a mean value of 3.7% at the deeper site MD98-2166 (Figure 4d).

As can be readily seen in Figures 2, 3 and 4, standard deviations, in each data set, integrate a few "anomalous" data points for which the batch A-batch B aliquot differences are negative, meaning that the reductively cleaned sub-samples show higher Mg/Ca ratios than their corresponding non-reductively cleaned aliquots. These anomalous values are particularly high in the G. menardii dataset from core MD98-2166, with one sample showing a non-reductively cleaned A-aliquot with 29% less Mg than the corresponding B-aliquot (Figure 4).

Our results also show that there is no apparent co-variance between relative Mg/Ca loss in the three planktonic species along each two cores (see panels (c) and (d) in Figures 2, 3 and4).

This suggests that most of the variance in A-B Mg/Ca differences is random, reflecting inhomogeneities between aliquots, and does not reflect any strong pre-conditioning of the samples (e.g. through evolution of dissolution conditions at the seafloor during the deposition of sedimentary material). Such random scattering around the mean values is also confirmed Shapiro-Wilk tests (i.e. null hypothesis of normal distribution with p>0.05), which indicate that Mg/Ca differences between paired A-B aliquots follow normal distributions.

Although we cannot reject the possibility that scattering and these anomalous negative values (i.e. Mg/Ca in non-reductively cleaned samples < Mg/Ca in reductively cleaned samples) result to some extent from analytical noise, this effect should be only marginal as there is no evidence of severe and rapid instable ICP-MS behavior during runs. Anomalous Mg/Ca paired-values are not associated, for instance, with anomalous values of other elements measured together with Mg and Ca. Thus, we believe that variances of several percent around the averaged Mg/Ca differences chiefly result from inhomogeneity between aliquots from the same samples, despite our efforts to mix thoroughly crushed foraminifer tests before splitting them into paired-aliquots.

In order to avoid excessive loss of calcite material during the ultra-sonification of the cleaning protocols, foraminifer tests were only slightly crushed to open shell chambers but not finely ground (Barker et al., 2003). It is known that Mg is hetMnrogeneously distributed in the shells of the three species used in this study: G. ruber (Eggins et al., 2003;Sadekov et al., 2009), P.

obliquiloculata (Sadekov et al., 2009;Steinhardt et al., 2015) and G. menardii (Sadekov et al., 2005). Large fragments of shell chambers may not allow for a thorough enough homogenization and could result in slightly different Mg/Ca composition between A and B aliquots. It is noteworthy that the two species of our data sets (namely P. obliquiloculata and G. menardii)

showing the highest variance in terms of Mg-cleaning loss and extreme anomalous negative values, are two species with complex structures characterized by a very heterogeneous distribution of Magnesium in the tests. P. obliquiloculata has a thick and dense outside cortex with lower Mg/Ca than the inside of the test (Steinhardt et al., 2015). G.menardii shows a heavy outer crust as well as thick keel (Regenberg et al., 2010), which are characterized by low Mg/Ca ratios (~ 2 mmol.mol -1 ) that contrast with the higher Mg/Ca (ranging from 4 to 6 mmol.mol -1 )

of the inner chamber calcite (Sadekov et al., 2005). G. ruber, on the other hand, is a bi-lamellar species: i.e. each new layer of calcite is constituted by an inner, Mg-rich primary layer and an outer, Mg-poorer layer. Thus, contrary to G. mernardii and P. obliquiloculata -whose heterogeneity occurs at the test-scale -Mg heterogeneity in G. ruber occurs at the wall-scale, and breaking open the tests should not result in significant Mg/Ca differences between the large fragments. This may explain the lower standard deviation shown by Mg-difference between aliquots of G. ruber samples compared to the two other species.

We conclude from the above observations that Mg/Ca differences between paired A-B aliquots of a given species chiefly reflect the excess-dissolution of Mg-rich layers induced by the reductive cleaning procedure (Barker et al., 2003;Rosenthal et al., 2004), with no clear impact from the removal of oxides (i.e. no relationship between Mg and Mn relative losses). Our data show that a rather large scattering can be superimposed on the relative Mg/Ca offset between the cleaning methods. We attribute the particularly high scattering to inhomogeneities in Mg distribution within the foraminifer test (i.e. and, thus, between the fragments), which randomly affect the Mg difference between A and B aliquots, leading (i) to a higher scattering in G.

menardii and P. obliquiloculata datasets than in G. ruber dataset and (ii) resulting in odd, negative offsets for some samples (i.e. Mg/Ca in A aliquots < Mg/Ca in B aliquots).

What controls the relative loss of Mg during reductive cleaning?

As indicated above, the reductive cleaning protocol results in a systematic and important relative loss of manganese in Batch-B aliquots compared to batch-A aliquots. The relative losses of Mn/Ca range from 56% for G. menardii in core MD98-2166 (mean Mn/Ca dropping from 0.22 mmol.mol -1 to 0.1 mmol.mol -1 when adding the reductive step) up to 72% for P.

obliquiloculata in core MD98-2166 (the mean Mn/Ca dropping from 0.21 mmol.mol -1 to 0.07 mmol.mol -1 ). The Mn losses average 62% when computing together data from the three species at both sites. Such an important Mn loss traduces the efficient removal of Mn-rich oxides when using the reductive cleaning protocol. The fact that reductively cleaned samples still contain a small but measurable amount of Mn suggests that this element is also present in the calcite lattice. Whether this Mn is located in the foraminifer test or in a diagenetic Mn-rich carbonate phases (Boyle, 1983;Pena et al., 2005) is beyond the scope of this paper. Importantly, there is no relation between Mn and Mg contents and there is no relationship between the relative loss of Mn and that of Mg, suggesting that -in our two cores -the cleaning-related loss of Mg does not chiefly result from the removal of Mn-rich oxides that may have contained Mg among other metals. Thus, our data confirm that the decrease of Mg/Ca ratios associated with the reductive cleaning likely results from cleaning-induced, calcite dissolution, which preferentially removes the more labile and Mg-richer parts of foraminifer tests (Barker et al., 2003;Rosenthal et al., 2004). Recently published MEB images provide direct evidence of such cleaning-induced dissolution effects on foraminiferal microstructures (e.g., Johnstone et al., 2016;Peral et al., 2018). synthetic plot provides the best graphic solution to summarize the distribution and relative contribution of each species and core to the overall estimate of Mg/Ca loss (Figure 5). The slope obtained through a simple linear regression approach has been used in many previous studies to provide an estimate of the average Mg differences between A and B aliquots (Barker et al., 2003;Rosenthal et al., 2004;Yu et al., 2007). When taking into account the three species together, at the two sites, the overall Mg-loss is ~ 5.7% (i.e. slope 0.943 of the green regression line, Figure 5). This loss is much smaller than the 15% loss found by Rosenthal et al., (2004) from their multiplanktonic foraminifera study. It is noteworthy that their estimate of a 15% decrease due to reductive cleaning was obtained from a linear regression approach without forcing the intercept to zero, whereas we chose to force the intercept to zero in Figure 5. Forcing the regression to zero assumes that the linear tendency we observe over the 1.5-4.5 mmol/mol Mg/Ca range of our dataset can be strictly extrapolated down to a zero Mg loss in a pure calcite (i.e. no Ca substitution). By forcing the intercept to zero, the average loss we obtain through this approach is similar to that obtained by computing a simple arithmetic mean of the whole dataset. We used published data from Rosenthal et al. (2004, Figure 5 therein), and run a linear regression with an intercept forced to 0. The re-evaluated slope indicates a reductive Mg loss of 6% for the multi-species, foraminifer Mg/Ca database of Rosenthal et al. (2004); a result fully consistent with our own. This suggests that differences between published Mg-loss may not only result from different protocols or different species or assemblages being analyzed, but may also depend upon the approach used to estimate the average Mg/Ca decrease.

Cleaning effect on

Species-dependence of Mg loss during reductive cleaning

Merging all the data together provides an estimate of cleaning Mg/Ca loss (~ 5.7%) that is coherent with the revisited Mg/Ca loss (6%) using the database of Rosenthal et al. (2004). Such multi-species estimates of cleaning loss mask, however, species-dependent differences, which appear statistically significant when applying Student or Welch's t-tests.

In order to highlight the potential species-dependent differences in Mg/Ca resulting from the cleaning process, we focus on the data from the shallower of our two cores, MD98-2165, which provides well-preserved planktonic tests. Among the three species we studied, P.

obliquiloculata shows the smallest average loss of Mg (~ 4.4%, Figure 3c). This average Mgloss is significantly higher than the 1.8% Mg-loss found by Xu et al., (2008) and the 0% loss estimated by Johnstone et al., (2016) for this species. But the average loss we found is slightly smaller than the ~ 6% loss published by Fan et al. (2018) for P. obliquiloculata tests picked from two Makassar Strait cores (Indonesian Archipelago) and retrieved at shallower water depths (1185m and 1984m) than MD98-2165 (2100m).

Our data indicate that G. menardii shells lose almost twice as much Mg (8.5%) than P.

obliquiloculata (4.4%) in core MD98-2165 when a reductive step is added to the cleaning protocol (Figure 4c). The third species studied, G. ruber, shows an intermediate, cleaningrelated Mg loss, which averages 6.2 %. This result is consistent with the 6.6% decrease observed by Xu et al., (2008) using material from a nearby sediment core retrieved at a similar water depth (1783m) in the Indonesian Archipelago. Interestingly, Xu et al. (2008) and the present study obtain the same Mg-loss for G. ruber (~ 6%), whereas the two studies differ significantly with regard to the P. obliquiloculata Mg-loss (see above).

The glossy P. obliquiloculata test is one of the most dissolution-resistant among planktonic foraminifers (Berger, 1967;Johnstone et al., 2016). Increasing the substitution of Ca by Mg is known to degrade calcite crystallinity and increase its sensitivity to dissolution (Bassinot et al., 2004;Nouet & Bassinot, 2007). We believe that the smaller decrease of Mg during the reductive cleaning (i.e. higher resistance to dissolution) shown by P. obliquiloculata compared to the two other species, results from the lower Mg content of its calcite test (with Mg/Ca ranging from ~ 2.0 mmol.mol -1 to 2.7 mmol.mol -1 in core MD98-2165).

Interestingly, G. menardii not only losses twice as much Mg (8.5%) than P. obliquiloculata (4.4%) when cleaned with the reductive protocol, but it also losses more Mg than G. ruber (6.2%). This higher relative loss of Mg from G. menardii seems to be at odd with the fact that G. ruber is considered a dissolution-sensitive planktonic species, whereas G. menardii is considered a dissolution-resistant one (Berger, 1967). However, it is important to recall that the capacity of G. menardii to resist dissolution is chiefly due to the presence of a thick and low-Mg keel (Brown & Elderfield, 1996;Sadekov et al., 2005). This keel can be preserved in sediments as easily identifiable, crescent-shaped fragments, even when the chamber walls have long been dissolved and G. menardii tests reduced to fragments (Mekik & François, 2006;Mekik et al., 2007). If not for this resistant keel, G. menardii should not be considered a dissolution-resistant species since its walls, which are two to three times richer in Mg than the keel (Sadekov et al., 2005), can show very early signs of dissolution (i.e. holes starting to develop early in the external walls of the last foraminifer chambers and rapidly increasing in size with proceeding dissolution). The early impact of dissolution on its chamber walls, associated with the resistance of its keel make G. menardii an interesting candidate for developing semi-quantitative proxies of dissolution intensity at the seafloor (Mekik & François, 2006). We believe that it is this capacity of chamber walls to dissolve rapidly that explains why G. menardii tests from core MD98-2165 losses more Mg than P. obliquiloculata and even more than G. ruber when subject to the reductive cleaning.

The effect of in situ dissolution in modulating Mg/Ca cleaning reduction

When considering separately Mg/Ca data from the two cores (Figure 5), planktonic foraminifers picked from the shallow core MD98-2195 show a slightly higher Mg-loss during reductive cleaning (slope of 0.937, i.e. loss of ~ 6.3‰,) than foraminifers from the deeper MD98-2166 core (slope of 0.952, i.e. ~ 4.8% loss). The difference is small but consistent with the hypothesis that foraminifers having undergone in situ dissolution are likely to lose less Mg during cleaning process than better-preserved tests (Barker et al., 2003). However, the impact of in situ dissolution is actually a bit more complex than just a reduced cleaning loss of Mg with increasing water depth of deposition. Our data suggest, here again, that there are some speciesrelated differences in the way in situ dissolution impacts Mg loss during cleaning.

As just seen above, G. menardii tests from the shallow core MD98-2165 show a significantly higher cleaning-related Mg/Ca decrease (8.5%) than the other two species (6.2% for G. ruber, 4.4% for P. obliquiloculata). In the deeper core MD98-2166, however, G. menardii exhibits a similar or lower cleaning-related Mg/Ca decrease (3.7%) than the other two species (4.1% for G. ruber, 6.3% for P. obliquiloculata). We interpret the strong difference in G. menardii Contrary to what is observed for the two other species, P. obliquiloculata exhibits a stronger, cleaning-induced Mg/Ca decrease in the deeper core MD98-2166 (6.3%), than in the shallower core MD98-2165 (4.4%) (Figure 3). This difference is seen as statistically robust when running a t-test. At present we have no explanation for such a puzzling observation and further studies, including MEB observation and elemental mapping in the calcite tests, will be mandatory to decipher why more intensively dissolved P. obliquiloculata specimens apparently lose more Mg during reductive cleaning than tests from the shallower, less dissolved core.

Including or not the reductive step in the cleaning method?

There is no doubt that using a reductive step during cleaning is necessary for those specific elements that are abundant in oxide coatings (e.g. Cd, U and Ba; Boyle & Keigwin, 1985;Yu et al., 2007), as well as, apparently, for Mg/Ca analyses performed on samples particularly affected by oxide contamination, such in the Panama basin (Pena et al., 2005(Pena et al., , 2008)). For most of non-problematic regions, like our study area in the Indonesian Archipelago, adding a reductive step does not improve Mg/Ca measurements. We observed no relationship between cleaning-related loss of Mg and Mn, and there are no outliers showing anomalously high Mg/Ca ratios associated to high Mn/Ca values in the non-reductively cleaned (A) samples. This suggests that the amount of Mg associated to oxides is negligible compared to the amount of Mg that is contained in the foraminiferal calcite, an observation already made by Barker et al. (2003). Reductive cleaning does not apparently improve Mg/Ca results with respect to the removal of oxides, but it significantly reduces the Mg/Ca ratios through the preferential dissolution of Mg-rich parts of the foraminiferal tests. In addition, as can be readily seen from Figures 2, 3 and 4, anomalous Mg/Ca values are more often found in the reductively cleaned datasets, an observation which could reflect the fact that the reductive cleaning adds more manipulations (i.e. increased likelihood of contamination) and is more difficult to conduct properly (i.e. it requires a precise timing of the reductive step to avoid variability in dissolutioninduced changes). Thus, we believe that the reductive cleaning should be used only in sediments with severe oxide coating of the particles and/or when the paucity of planktonic foraminifers makes it impossible to conduct separate analyses for Mg/Ca and other elements that are significantly contaminated by oxides (e.g. Cd, Ba,..). If one is interested chiefly in Mg/Ca, we believe that it is best to use the non-reductive approach for cleaning the foraminiferal tests. In cases when the reductive step must be included, attention should be paid on Mg/Ca results from species having strongly non-homogeneous distribution of Mg. Our results show, for instance, that in order to obtain reliable Mg/Ca-based temperatures from G. menardii, it is mandatory to work on well-preserved samples and the use of reductive step during cleaning should be better avoided.

Correcting for the Mg/Ca offsets induced by the cleaning methods

When dealing with spatial temperature distribution and trying to map temperature anomalies at a given period of the past (e.g. Last Glacial Maximum), it is crucial to insure the compilation of the most coherent temperature estimates possible. In such cases, efforts should be made to use as much as possible Mg/Ca data obtained with the same cleaning protocol. However, to date, no agreement has been reached regarding the cleaning protocol for Mg/Ca analyses. Some laboratories add a reductive step, while others do not include it. As a consequence, when comparing data from different laboratories or dealing with the construction of Mg/Ca-Tº databases, one has to cope with the potential problem of mixing inhomogeneous Mg/Ca datasets (e.g. Xu et al., 2010;Fan et al., 2018;Wang et al., 2018). This problem does not only concern direct Mg/Ca-Tº reconstruction. Mg/Ca-derived temperatures are also used, for instance, to reconstruct other paleoceanography proxies, such as the δ 18 O of seawater (LeGrande & Schmidt, 2006;Thirumalai et al., 2016), or to study the thermocline structure based on the Mg/Ca-Tº gradient between shallow and deep-dwelling species such as G. ruber and P. obliquiloculata (e.g., Holbourn et al., 2011;Fan et al., 2018;Hollstein et al., 2018;Wang et al., 2018). In all these cases, cleaning-induced differences in Mg/Ca can bias the interpretations. It is therefore crucial to decipher how we can minimize cleaning-induced Mg/Ca differences between datasets.

Our results and those from previous studies show that simple linear corrections can be applied.

In Figure 6 (a-b), we compare temperature records obtained in core MD98-2165 using the initial Mg/Ca results for G. ruber (panel a) and P. obliquiloculata (panel b) converted to temperatures using the empirical equation Mg/Ca = 0.38 exp (0.09*Tº) for G. ruber and Mg/Ca = 0.328 exp (0.09*Tº) for P. obliquiloculata (Anand et al., 2003). Without any correction for cleaning effects (upper curves), the average temperature difference between reductively cleaned and non-reductively cleaned samples is ~0.7 ºC for G. ruber and ~ 0.5°C for P. obliquiloculata. As expected these differences cancel out (lower curves) when correcting reductively cleaned

Mg/Ca values with species-specific offsets (i.e. +6.2% for G. ruber and +4.1% for P.

obliquiloculata) prior to temperature conversion. Using the same multi-species, average offset factor (5.7%, figure 5) to correct Mg/Ca data from both G. ruber and P. obliquiloculata would have resulted in slightly under-corrected G. ruber Mg/Ca-SST and over-corrected P.

obliquiloculata Mg/Ca-thermocline T°, thus reducing artificially the vertical temperature gradients that can be deduced from comparing the temperatures obtained with the two species.

Yet, using a common, multi-species correction factor does not modify the downcore evolution of the G. ruber and P. obliquiloculata temperature records, and does not affect, therefore, the interpretation of the relative evolution of vertical temperature gradient at site MD98-2165.

The necessity for as precise as possible quantitative estimates of temperatures is linked to our interest in comparing those temperatures between different sites of the same region. In order to illustrate potential misinterpretations that may result from cleaning issues and the interest of correction for these effects, we compared Mg/Ca-SST that we obtained from G. ruber in the shallow core MD98-2165, with G. ruber Mg/Ca-SST obtained from nearby core SO139-74KL (Wang et al., 2018) located to the South off Sumatra (Figure 1). Mg/Ca ratios in SO139-74KL

were measured from G. ruber in the 250-350 µm size fraction and the cleaning method included a reductive step (Wang et al., 2018). Mg/Ca ratios were converted to SST by using, for both records, the same empirical equation for G. ruber: Mg/Ca = 0.38 exp (0.09*Tº) (Anand et al., 2003). As can be seen from Figure 6c, Mg/Ca-SST obtained using the same cleaning method (i.e. with reductive step) indicate roughly comparable SST between the two cores during the early onset of deglaciation (20 -18 ka Cal BP) and the early Holocene (starting around 10ka), with a strong temperature gradient developing during the deglaciation itself. However, when we compare Mg/Ca-SST obtained using different cleaning methods (i.e. using MD98-2165 data obtained without the reductive step), the data suggest that a lateral temperature gradient always existed between the two sites since the LGM and only collapsed sometimes during the mid-Holocene (Figure 6d). Thus, paleo-interpretation is clearly affected by the cleaning effects on Mg/Ca records. We applied the G. ruber +6.2% correction to reductively-cleaned Mg/Ca data from core SO139-74KL (Figure 6e). After correction, the evolution of both Mg/Ca-SST records is fully coherent with the interpretation resulting the comparison of Mg/Ca-SST obtained with the same cleaning technique (Figure 6 (c). This example demonstrates the necessity to adjust the cleaning-offset when precise, absolute Mg/Ca-based water temperatures are needed (i.e.

vertical or lateral gradients) for paleo-interpretation.

Although a multi-species correction factor (i.e. 5.7%, see figure 5) is practical to use to minimize cleaning-offsets, we suggest using species-dependent correction factors when possible (i.e. 6.2% for well-preserved samples of G. ruber and 4.1 % for well-preserved samples of P. obliquiloculata). For the more poorly-preserved samples of G. ruber (such as the case in MD98-2166), a smaller correction factor of 4% should be more appropriate. For the poorlypreserved samples of P. obliquiloculata and G. menardii, it would be inappropriate to correct the cleaning-offset by a universal factor, as their large variance and weak correlation between Mg/Ca ratios cleaned with and without the reductive step; in fact, it is as well as pointless to correct them because they have lost the ability to reflect the water temperature due to the altered Mg/Ca ratios caused by the in situ dissolution. For those other species not mentioned in this study, the multi-species correction factor of 5.7% could be employed.

Conclusion

In this study, we compared the Mg/Ca ratios obtained using two cleaning methods on three planktonic species, G. ruber, P. obliquiloculata and G. menardii, picked from two sediment cores retrieved at different water depths. The Mg/Ca records cover the last ~21ka Cal BP for core MD98-2165 (2100m) and ~35ka Cal BP for core MD98-2166 (3875m).

Our results show that adding a reductive step decreases the Mg/Ca ratio for all three species.

The overall loss averages 6%. In the shallow, well-preserved Core MD98-2165, T-tests show that differences between mean Mg% losses in the three species are statistically robust, indicating that the cleaning loss is species-specific to some extent. The cleaning-related differences in Mg/Ca of G. ruber and G. menardii are higher in the shallower core MD98-2165 than in the deeper core MD98-2166, suggesting that foraminifera tests which have experienced intense dissolution on the sea-floor are less prone to lose further Mg during the reductive cleaning. P. obliquiloculata shows a very puzzling behavior, yet to be explained, with a higher Mg loss due to cleaning in the deeper site (6.3%) than in the shallower site (4.4%).

Although G. ruber is a species sensitive to dissolution, there are no strong differences in Mg/Ca between the shallow and the deep sites. Assuming that surface conditions were always similar at the two sites, this observation suggests that the dissolution does not strongly modify the initial Mg/Ca-SST. Yet, our data indicate that preservation state has a pre-conditioning effect on the capacity of G. ruber to lose Mg during the cleaning procedure, the average Mg/Ca cleaningloss reaching only 4.1% at the deeper site MD98-2166, while it is 6.2% at the shallowest site.

Well-preserved G. menardii shells retrieved in core MD98-2165 shows the strongest cleaninginduced Mg/Ca decrease of our study (8.5%), which is likely due to the fragile nature of the wall calcite that tends to dissolve easily. G. menardii specimens from Core MD98-2166, which have experienced some dissolution at the seafloor, show a marked reduction in cleaning-related Mg/Ca loss (3.7%) compared to specimens from the shallow core. We interpret this result as being indicative of the strong resistance to dissolution of the remaining outer crust and keel, characterized by low Mg/Ca composition (Sadekov et al., 2005). This suggests that wellpreserved G. menardii specimens (retrieved from shallow cores, always above the lysocline) are mandatory for reconstructing unbiased Mg/Ca-based thermocline temperature, and nonreductive cleaning procedure is clearly preferable for this species to prevent partial dissolution.

For all species studied, reductive cleaning results in a clear Mg/Ca decrease compared to nonreductively cleaned values. However, there is no sign of significant changes in the relative Mg/Ca variations along a given core and for a given species when comparing the curves obtained by the two cleaning methods. In other words, relative seawater temperature changes reconstructed along sediment records (e.g. the temperature change between LGM and Holocene)

are not affected by the choice of the cleaning method.

It is the authors' contention that non-reductive cleaning approach should be preferred for building inter-laboratory homogeneous datasets of Mg/Ca-derived temperatures, whereas the reductive step should be only included when severe oxide coating is observed.

Introduction

The Inter-Tropical Convergence Zone (ITCZ) is characterized by intense evaporation and atmospheric convection processes, which result in heavy rainfalls (Schneider et al., 2014).

Along the course of the year, the ITCZ migrates towards the warmer (i.e. summer) hemisphere in response to seasonal and latitudinal changes in insolation controlled by the orbital parameters of the Earth (Haug et al., 2001;Schneider et al., 2014). This latitudinal displacement of the ITZC characterizes the areas belonging to the Global Monsoon system (Wang et al., 2005). On longer time scales, orbitally-driven changes in seasonal insolation continue to play a key role in low-latitude hydrology, as can be readily seen from the occurrence of orbital periodicities (more particularly precession) in low-latitude, paleo-hydroclimate records (Ayliffe et al., 2013;Carolin et al., 2013;Cheng et al., 2016;Denniston et al., 2017). However, changes in insolation are not the only mechanisms controlling inter-hemispheric temperature gradients (Toggweiler & Lea, 2010). Those temperature gradients also reflect complex internal feedback mechanisms, which involve the redistribution of heat by oceanic circulation and its interactions with ice sheet development and dynamics, leading to the so-called Bipolar Seesaw (Buizert et al., 2018;Denton et al., 2010;Toggweiler & Lea, 2010;Barker et al., 2011).

Paleo-records of sea-surface δ 18 O reconstructed from marine sediments ( Levi et al., 2007;Gibbons et al., 2014;Mohtadi et al., 2014) and of continental precipitations reconstructed from cave stalagmites (Ayliffe et al., 2013;Carolin et al., 2013;Denniston et al., 2017;Krause et al., 2019) have revealed the occurrence of abrupt hydrologic changes in the IPWP over the last glacial period and during the last deglaciation. The most recent of these rapid, hydroclimate changes have been coeval with Heinrich Stadial 1 (HS1) and the Younger Dryas (YD), two cold events within the Last Glacial and the early deglaciation, which were triggered by the rapid destabilization and surge episodes of Northern Hemisphere ice-sheets. Paleoclimatic studies suggest that mechanisms involved in the tele-connection between high, northern latitude climatic events and tropical precipitation changes operated via the disturbance of the ITCZ ( Levi et al., 2007;Ayliffe et al., 2013;Carolin et al., 2013;Gibbons et al., 2014;Mohtadi et al., 2014;Krause et al., 2019). Denniston et al. (2017) convincingly strengthened such a hypothesis by showing that abrupt changes in precipitations recorded in stalagmite δ 18 O records are anti-phased between the northern and southern IPWP, suggesting that they result from the southward shift of ITCZ mean position during HS1 and YD cold events.

Transient modeling of hydrographic variations across the last deglaciation suggests that past changes in surface ocean d 18 O in the tropics were strongly controlled by the average position and dynamics of the ITCZ, reflecting the evolution of the mean state of inter-hemispheric temperature gradients (Gibbons et al., 2014). This means that long d 18 O record could be used to explore the dynamics of the ITCZ at orbital time scales and at key periods of the past characterized by different boundary conditions (insolation, atmospheric pCO 2 , ice volume). So far, such issues have not been properly addressed, however, owing to the paucity of long records of surface ocean δ 18 O that are properly corrected for past changes in temperature and ice sheet volume.

In this study, we provide a high-resolution record of surface ocean d 18 O in the IPWP region.

This record has been derived by combining δ 18 O data and Mg/Ca-temperatures obtained on the planktonic foraminifer Globigerinoides ruber picked from sedimentary core MD98-2165. This record allows to study past hydroclimate changes in the IPWP region and explore the mechanisms that control them at different timescales.

Modern settings

The region comprising North Australian and the South-East Asian region is characterized by seasonal shifts of prevailing winds during boreal summer (June, July and August -JJA) and winter (December, January and February -DJF) accompanied by contrasted dry and rainy seasons, which is known as the Australian-Asian Monsoon system. The coupling between the South-East Asian monsoon and the Malaysian-Australian monsoon sub-systems rests upon the latitudinal meandering of the ITCZ. This maximum rainfall belt typically moves from ~10°S to ~30°N (Supplementary Figure S1) along the course of the seasons, migrating towards the warmer (i.e. summer) hemisphere in response to latitudinal changes in insolation controlled by the relative position of the Sun in the sky.

Sediment core MD98-2165 (9.65 °S, 118.34 °E) is located in the south of equator between Australia and Indonesian (Figure 1). In this region, during austral summer season (DJF), relatively warm and moist winds blow from northwest, bringing heavy rainfall, with the ITCZ reaching its southernmost position (Supplementary Figure S1). Conditions reverse during the austral winter season (JJA), with relatively cold and dry winds blowing from southeast and rainfall reaching its lowest value of the year at the site of MD98-2165. The ITCZ is at its northernmost position during this period, locked north of the equator. Thus, at the site of core MD98-2165, about ~70% of total year rainfall occur from December to March, with precipitation rate reaching 11-12 mm/day. By contrast, during July to August, the precipitation rate is only 1-2mm/day in average (Supplementary Figure S2).

The main local current prevailing at the site of core MD98-2165 is the Indonesian Throughflow (ITF) (Figure 1). The ITF flows from the Pacific Ocean to the Indian Ocean through the Indonesian Archipelago. Modern observations indicate that over the course of the year, an average ~15 Sv (1 Sv = 10 6 m/s) of water are carried by the ITF to the Indian Ocean through three main exits: The Lombok strait, the Ombai strait and Timor passages) (Sprintall et al., 2009). The ITF conveys surface and sub-surface waters, down to a depth of about 700m meters, although due to shallow sills between some islands of the Indonesian Archipelago, only surface waters are carried through in many areas. Core MD98-2165 is located in the ITF outflow region, south of the Lombok strait (Figure 1). The ITF water is characterized by low salinities compared to the Indian Ocean water, resulting in a clear fresh water tongue at both surface and subsurface water depths, which extend from the eastern to the central Indian Ocean (Figure 2). (Visser et al., 2003). GeoB10069-3 (Gibbons et al., 2014); MD01-2378 (Xu et al., 2008); Flores Cave (Ayliffe et al., 2013); Borneo Cave (Meckler et al., 2012;Carolin et al., 2013). China Caves (Cheng et al., 2016). (Mohtadi et al., 2014). 2:SO139-74KL (X. Wang et al., 2018). 3: MD98-2165 (this study). 4:MD98-2162 (Visser et al., 2003). 5:

GeoB10069-3 (Gibbons et al., 2014). 6:MD01-2378 (Xu et al., 2008). (AMS) 14 C dates (Table 1), already published by Waelbroeck et al. (2006), suggest that the top anomalously high Mg/Ca-derived SST down to 70 cm, although the 14 C age at 55 cm is perfectly coherent with the depth-age profile. We focused the present study in the interval 90-3450 cm (interpolated age at 90 cm is ~ 1.2 ka).

Material and Methods

High resolution δ 18 O records of both planktonic and benthic species have been obtained for the complete core in previous studies and published elsewhere (Levi, 2003;Levi et al., 2007;Waelbroeck et al., 2006). Planktonic δ 18 O were measured on mono-species Globigerinoides ruber (250-315 um; white) (Levi et al., 2003;Levi et al., 2007), benthic δ 18 O from Cibicides wuellerstorfi (>150 µm) for the whole interval of 90 -3450 cm, except for the part of 1996 -2608 cm (corresponding to Termination II) where Hoeglundina elegans (>150 µm) species was used (Levi 2003;Waelbroeck et al., 2006).

Age Model

An initial age model for the upper section (55 to 774 cm) had been published based on accelerated mass spectrometry (AMS) 14 C dates measured from mono-specific G. ruber white (Waelbroeck et al., 2006). Calendar ages had been computed by using the 1998 marine calibration curve with a surface reservoir age of 400 years (Waelbroeck et al., 2006). Here, we use the latest MARINE13 age calibration curve (Reimer et al., 2013) to convert AMS 14 C dates to calendar age through Calib 7.10 software (Table 1). Marine reservoir ages were set by default to 405 years. The new generated calendar ages are consistent with previously published ages for 14 C dates younger than 16 ka, and ~200-500 years older for 14 C age in the range 16-26 ka.

For the older part of the core, the age model was established by aligning visually the benthic δ 18 O record of MD98-2165 to the global benthic δ 18 O stack (Lisiecki & Raymo, 2005;Lisiecki & Stern, 2016) using the AnalySeries software (Paillard et al., 1996). We used 5 tie points to align the records (Figure 3e). Depths were converted to age by linear interpolation between AMS 14 C calendar dates or oxygen stratigraphy tie points. Based on this age model, the 90 -3450 cm depth interval studied in the present paper covers the time interval from ~1.2 to 270 ka. The average sedimentation rate is ~13cm/ka.

Table 1. AMS 14 C dates and corresponding calendar ages of core MD98-2165 * 14 C dates were taken from Waelbroeck et al. (2006).

Mg/Ca-based SST reconstruction

The Mg/Ca measurements of G. ruber from the 90 -800 cm interval are from the study of Levi et al., (2007). Here, we extended the G. ruber Mg/Ca dataset downward to 3450 cm. About 30 tests of G. ruber (white) were picked every 4 to 12cm in the 110-3450 cm interval from the 250-355 µm size fraction. As was done by Levi et al., (2007), samples were cleaned without a reductive step following the method described in Barker et al. (2003) The δ 18 O cal of foraminiferal calcite is controlled by both the temperature (via fractionation mechanisms evidenced by the seminal work of [START_REF] Urey | The thermodynamic properties of isotopic substances[END_REF]) and by the δ 18 O sw of ambient seawater in which foraminifers calcify their shells [START_REF] Emiliani | Pleistocene temperatures[END_REF]Bemis et al., 1998) . We used the relative sea level variations established by Waelbroeck et al. (2002) to correct for global mean ocean δ 18 O changes. Other global sea-level records (i.e., Grant et al., 2012;Spratt & Lisiecki, 2015) were also used for testing the sensitivity of δ 18 O sw-ivc to different reconstruction of global ice-volume changes (i.e. sea-level) and their effects on ocean mean δ 18 O. We also used the benthic δ 18 O benthic record from Core MD98-2165 to simulate global δ 18 O changes. As we have no benthic temperature record for core MD98-2165, we assumed that changes in the δ 18 O benthic record were chiefly controlled by variations in ice volume. We know this assumption to be an 100 oversimplification, but it is our contention that past changes in bottom water temperature marginally affected δ 18 O benthic in Core MD98-2165. Indeed, between the Last Glacial Maximum (LGM; ~ 4‰) and the Holocene (~ 2.6‰), the amplitude of the δ 18 O benthic variation is ~ 1.4‰, whereas the global ice volume effect accounts for a ~ 1.2‰ change (estimate based on the 0.010‰m -1 relationship derived from pore water studies; (Adkins, 2013)). This suggests a limited temperature effect, accounting only for a 0.2‰ change over the last glacial termination.

Although such temperature effect may have noticeable impacts on the precise time-evolution of δ 18 O benthic reconstructed changes across major deglaciations (e.g., Lisiecki & Stern, 2016;Skinner & Shackleton, 2005), in the present study we neglected this effect and simply converted the δ 18 O benthic of MD98-2165 into a global, ice-volume related change in ocean δ 18 O by normalizing the last deglaciation to match a 1.2‰ glacial-interglacial change. Despite potential drawbacks of this approach, it is the benefit from combining global (benthic) δ 18 O effect and local, planktonic sea surface δ 18 O from the same core, thus avoiding any problem related to age model and improper alignment to global δ 18 O estimates. Interestingly, as can be seen in Figure 3c, the benthic δ 18 O record of MD98-2165 is perfectly in phase with the ice volume record of Waelbroeck et al. (2002), whereas it lags significantly during the two previous terminations.

This might result from a potential problem in age models, suggesting that using the benthic δ 18 O would be definitively a better strategy when correcting the planktonic δ 18 O for global effect.

Using bootstrap Monte Carlo simulations within toolkit PSU-Solver including the analytical, calibration and sampling uncertainties, we estimated the average uncertainty of δ 18 O sw-ivc to be ~ 0.2‰ (1s). In order to provide an as coherent and robust comparison of δ 18 O sw-ivc records available from the literature with data from MD98-2165, we retrieved the published Mg/Ca and δ 18 O data and (re-)estimated δ 18 O sw-ivc using the PSU-Solver toolkit with the same settings (i.e.

calibration equations and sea level curve) that were used for MD98-2165.

Results

G. ruber Mg/Ca-SST

G. ruber SST from the top interval studied (~1.2 ka) is ~ 27.8°C, similar to the modern annual mean SST of 28.2°C. The G. ruber Mg/Ca record exhibits typical glacial-interglacial oscillations over the last ~270 ka, with higher (colder) temperature during interglacials (glacials), and displaying strong similarities with the benthic δ 18 O record (Figure 3b). The Mg/Ca values range between 3.2 and 5.3 mmol/mol, equivalent to SST change comprised between 23.8 and 29.2 °C.

The MIS 8 was warmer than the other two, subsequent glacial periods (MIS 6 and MIS 2-4).

During Termination III (deglaciation between MIS 8 and MIS 7), the clear onset of SST increase occurred at ~ 250 ka, in phase with the benthic δ 18 O change and leading the planktonic In order to extract the local δ 18 O signal, the global sea level component (i.e., ice volume changes)

was subtracted from the planktonic δ 18 O sw record from core MD98-2165. The corrected δ 18 O sw- ivc records obtained by subtracting different sea level curves (i.e., Grant et al., 2012;Spratt & Lisiecki, 2016;Waelbroeck et al., 2002) or subtracting the benthic δ 18 O record from MD98-2165, are similar within ± 0.2‰ (Supplementary Figure S3). The only significant difference (difference > 0.2‰) is at Termination II when using the past sea level curve of Spratt (2016).

The sea level curve of Spratt (2016) shows a relatively large offset with the other two sea level curves (i.e., Grant et al., 2012;Waelbroeck et al., 2002) as well as normalized, benthic-δ 18 O record of core MD98-2165. The local δ 18 O sw-ivc curve obtained by subtracting the sea level record of Waelbroeck et al. (2002) is the closest to the record obtained by averaging the different approaches, including the approach obtained by subtracting the normalized, benthicδ 18 O of core MD98-2165. In the rest of the paper, the δ 18 O sw-ivc obtained by using sea level curve of Waelbroeck et al. (2002) will be used and is presented in the figures for comparison.

As shown in Figure 3d salinity. This relationship can be empirically described using linear regression equations. In the Indian Ocean, the slope of such a linear relationship is ~0.16‰/p.s.u (LeGrande & Schmidt, 2011). Past changes in evaporation/precipitation intensity and location, combined with changes in atmospheric and ocean circulation, likely modified those regional linear relationships between salinity and δ 18 O sw . Modeling suggested that the modern relationship also applied during the LGM in the northern Indian Ocean (Delaygue et al., 2001). However, it is unlikely that modern isotopic/salinity relationship could be confidently used back in time, and in particular over the entire last 270 ka. However, even if we cannot provide robust quantitative estimates for paleo-salinities, the δ 18 O sw-ivc record of core MD98-2165 can still be used to decipher relative changes in salinity resulting from changes in hydroclimate and lateral advection of surface water. In this study, we will interpret the higher (lower) δ 18 O sw-ivc values as representing relatively saltier (fresher) seawater.

Impact of ITF evolution on eastern Indian Ocean δ 18 O sw-ivc change

At present, at the core site of MD98-2165, the annual mean salinity is ~34‰ which is lower than the average salinity of the central Indian Ocean (>35‰) (Figure 2). This lower salinity is induced by the fresh water input from the north Indonesian Seas carried by the ITF. The freshening effect of ITF on the Indian Ocean can be readily seen from the modern salinity distribution map (Figure 2). Seasonal changes of the surface salinity show that changes in the strength and/or salinity of ITF surface water are the main drivers of salinity changes at the site of core MD98-2165. By analogy with modern, seasonal variations, we can make the hypothesis that past changes in ITF strength affected salinity variations at the site of core MD98-2165. For instance, if the ITF had been weaker in the past, saltier Indian Ocean waters could spread more easily over the site of Core MD98-2165, leading to higher salinity and higher δ 18 O sw . Thus, in addition to the potential local precipitation and evaporation impact, the salinity of surface waters carried by the ITF or/and the strength of ITF should be other potential drivers of the changes in past δ 18 O sw-ivc records observed in Core MD98-2165.

In order to disentangle the complex δ 18 O sw-ivc record of MD98-2165, and to be able to more efficiently separate evaporation/precipitation impacts from lateral advection, we compared our The nearest cave stalagmite δ 18 O record to Core MD98-2165 is located in the Island of Flores (Ayliffe et al., 2013) (Figure 1). It is a much smaller record, though, covering only the last 30 ka. As can be seen from Figure 5, The Borneo and Flores records share common features at the glacial-interglacial timescale with high δ 18 O during the last glacial periods becoming lighter through the deglaciation and into the Holocene. One cannot exclude the possibility that a part of that common, glacial-interglacial signal could result from regional changes in precipitation.

However, the average amplitude of the δ 18 O change from the LGM to the Holocene (about 1 ‰) in both records suggests that this decreasing trend in d18O chiefly reflects the global change in ocean δ 18 O, thus impacting the isotopic composition of the water vapor that ultimately precipitates over Borneo and Flores. At short, millennial time scale, however, there are clear out-of-phase relationships between the Borneo and the Flores records (Figure 5). Noticeably, two anomalous enrichments can be observed in the MD98-2165 δ 18 O sw-ivc record during the last termination. These events appear to be coeval (within the limits of our age model) to cold YD and HS1 events in the northern Hemisphere, suggesting that these events triggered increased et al., 2013;Meckler et al., 2012). (b) Green curve is benthic δ 18 O of core MD98-2165 (Levi, 2003;Waelbroeck et al., 2006); black curve is global benthic δ 18 O stack, 0-150 ka is from LS16 (Lisiecki & Stern, 2016) and 150-270 ka from LR04 (Lisiecki & Raymo, 2005). (Waelbroeck et al., 2002). (d) Borneo cave δ 18 O record ( Carolin et al., 2013). (e) Flores cave δ 18 O record (Ayliffe et al., 2013). (f) Comparison of δ 18 O sw-ivc between MD98-2165 and other sediment cores. Although there are high similarities, there are also obvious discrepancies between MD98-2165 δ 18 O sw-ivc and Borneo cave δ 18 O records. For example, the increase of δ 18 O sw-ivc in MD98-2165 at ~ 60ka is comparable with the earlier increase observed across Termination II (Figure 4a). In comparison, in the Borneo cave δ 18 O, the enrichment at ~60ka is substantially weaker than the one occurring during Termination II. This observation either suggests that the coupling was not always strong between both sites, suggesting either a modulation by ITF intensity and/or intensified evaporation/precipitation in the vicinity of Core MD98-2165 (with a relatively lesser influence of evaporation/precipitation over the Borneo area). Similar discrepancies can be observed between the Borneo cave δ 18 O record and other cores located in the main path of the ITF (Figure 5 and6). As shown in Figure 5, the δ 18 O sw-ivc at core site of MD98-2162 located in the upstream region of ITF, within the Indonesian Archipelago, is lower than that in the ITF downstream region of cores MD98-2165, GeoB10069-3 and MD01-2378 during the interval 25 -16ka. This can be attributed to the strong "buffering" effect of high salinity Indian Ocean water due to relatively weak ITF during this period (Fan et al., 2018). Moreover, MD01-2378

shows no hint of change in δ 18 O sw-ivc during YD and HS1 (Xu et al., 2008), and the value of δ 18 O sw-ivc is generally higher than that of MD98-2165 during the B/A period. Core MD01-2378 is located in one of the main passages (Timor sea, Figure 1) for ITF into the Indian Ocean (Sprintall et al., 2009). And the position of core MD01-2378 is more close to the 35 p.s.u.

isohaline (Figure 2). This 35 p.s.u. isohaline should be controlled by the strength of ITF; weak ITF would allow the 35 p.s.u. isohaline shift northeastward. Therefore, the still enrichment of δ 18 O sw-ivc during B/A and as well as the early Holocene at the core site of MD01-2378 probably indicate weak ITF transport in the Timor passage, which place the core MD01-2378 in the control of high salinity Indian Ocean water from the northwest of Australia. At ~60 ka event, however, the δ 18 O sw-ivc was lower at core site of MD01-2378 than that in MD98-2165, suggesting relatively higher SSS at core site of MD98-2165. As aforementioned, Borneo cave δ 18 O show moderate increase at ~60ka compared to the period of Termination II (Figure 4).

This means that the moderate increased salinity in ITF water from the north cannot fully explain the anomalous higher δ 18 O sw-ivc at core site MD98-2165. In comparison, same degree increase in δ 18 O sw-ivc at ~60 ka is observed in the core SO139-74KL (Figure 4b). This results implies that high salinity Indian Ocean water from the western side reaching to the core MD98-2165, and this could explain the anomalous higher SSS at core site of MD98-2165. Similar condition is also observed during Termination II. As shown in Figure 6, δ 18 O sw-ivc records from core MD98-2162, MD98-2165 and MD01-2378 co-vary with the Borneo cave records during the period of 140 -110ka, supporting the ITF coupling the δ 18 O sw-ivc change in the path of ITF from Indonesian seas to the Indian Ocean. However, the δ 18 O sw-ivc value at core site MD98-2165 was higher than the sites of core MD98-2162 and MD01-2378 during the 132 -140 ka, and again similar with the value at site of SO139-74KL. This suggest that during the 132 -140 ka, the ITF effect was weak at core site MD98-2165, resulting saltier Indian Ocean water invade towards the ward of Indonesian seas, leading to the anomalously higher δ 18 O sw-ivc . This anomalously higher δ 18 O sw-ivc during the period of 132 -140 ka is responsible for forming an obvious increasing trend during MIS 6 which is not evident in the Borneo cave δ 18 O curve. All these three periods (16-24ka, 60ka and 132-140ka) mentioned above with anomalously higher δ 18 O sw-ivc at core site MD98-2165 compared to the ITF water occurred at periods with low stand sea level. The direct connection for the ITF from the Indonesian seas to the core MD98-2165 is via the Lombok Strait. The Lombok Strait is only ~300 m deep with a width of ~ 200km (Gordon, 2005). The mean ITF transport through Lombok Strait is ~ 2.6sv of total 15sv (Sprintall et al., 2009). Therefore, lower sea level may reduce the ITF transport through Lombok Strait, causing the higher salinity Indian Ocean water into the core MD98-2165 and leading to the anomalously higher δ 18 O sw-ivc .

Based on above discussion, we conclude that high similarity between Borneo cave δ 18 O and MD98-2165 δ 18 O sw-ivc is induced by the ITF linkage between the equatorial Indonesian Seas and the eastern Indian Ocean. Reconstructed past δ 18 O sw-ivc records from core MD98-2165 is paced by the change in δ 18 O of ITF seawater from the Indonesian Seas, which is affected by the change of rainfall in equatorial Indonesian Seas. The amplitude of MD98-2165 δ 18 O sw-ivc is modulated by the strength of ITF, weak ITF would lead to anomalously higher δ 18 O sw-ivc by allowing high salinity Indian Ocean seawater into the ward of Indonesian Seas.

AMOC forcing on hydroclimate changes in the IPWP

The large abrupt δ 18 O sw-ivc enrichments observed in ocean sediment core MD98-2165 as well as in the Borneo cave δ 18 O all occurred during transition periods when the sea surface temperature became warmer (e.g., at Terminations) and substantial ice sheet melting took place in high latitudes (Figures 3 and4). In contrast, the most depleted δ 18 O sw-ivc events occurred during cooling SST transitions (e.g., at ~230 ka, ~ 190 ka and ~70 ka), during periods coeval with substantial ice sheet accumulation (Figure 3 and4). These observations suggest the potential link between changes in the Indonesian seas hydrology, on the one hand, and processes resulting from the waxing and waning of high-latitude ice sheets.

There exist rather good correlations between MD98-2165 δ 18 O sw-ivc curve and 231 Pa/ 230 Th record from North Atlantic Ocean, a proxy of the strength of AMOC which is related to the dynamic change of ice sheet (Böhm et al., 2015). We find that episodes of weak AMOC are correlated to the increase of δ 18 O sw-ivc of MD98-2165, especially during T-II (~ 130ka), the MIS4/3 transition (~ 60ka) and T-1 (at ~16ka and 12ka) (Figure 7a). Weak AMOC can reduce the northward ocean heat transport, cooling the Northern Hemisphere and warming the Southern Hemisphere (Buizert et al., 2018), which change the interhemispheric thermal gradient. As ITCZ tends to move towards the warmer Hemisphere (Haug et al., 2001;Schneider et al., 2014), a weaker AMOC would result in the southward shift of the average ITCZ position.

The Borneo cave δ 18 O record indicates that during periods of weak AMOC (such as during Termination II at ~130 ka) the amount of rainfall around the equatorial Indonesian seas decreased. The out-of-phase relationships between Borneo and Flores records that result from the ITF dynamics (Denniston et al., 2017) suggest that at the location of our core MD98-2165

(nearer to Flores than to Borneo), in situ summer season rainfall likely increased in response to the southward shift of ITCZ (Figure 5). But, as already discussed in section 5.2, in situ seasonal rainfall changes at core site of MD98-2165 were probably not strong enough to affect the local SSS changes, and the δ 18 O sw-ivc of MD98-2165 rather reflect the relative SSS changes from upstream ITF in the Indonesian seas. The increased δ 18 O sw-ivc of MD98-2165 during periods of weak AMOC provide evidence that although seasonal rainfall change may vary spatially in the Indonesian area -as revealed by cave δ 18 O records (Ayliffe et al., 2013;Carolin et al., 2013;Denniston et al., 2017;Krause et al., 2019) -the average rainfall generally decreased as a result of southward shifts of ITCZ. In other words, a weak AMOC would promote dryer conditions over the Indonesian seas, leading to increased SSS of surface Indonesian seawaters and, therefore, the injection of saltier waters into the Indian Ocean.

Based on the fact that persistent good correlation between δ 18 O sw-ivc of MD98-2165, Borneo record (Figure 7). There is, however, a very strong relationship between the China cave δ 18 O and northern hemisphere summer insolation changes at orbital timescale (Cheng et al., 2016; Figure 7b). These two observations may suggest that past changes in the China monsoon were mainly driven by direct insolation and were marginally affected by changes in interhemispheric temperature gradient that were not forced by insolation (i.e. changes in the bi-polar seesaw and global redistribution of heat by oceanic circulation).

If such a hypothesis is valid, then a stronger correlation to the Indonesia area δ 18 O records would be seen after subtracting insolation-driven variations from the China cave δ 18 O record. (Cheng et al., 2016) and Borneo cave stalagmite δ 18 O (Carolin et al., 2013;Meckler et al., 2012), Sub-AM means the suborbital variability of the AM in (b) obtained by subtracting the orbital-scale variations (21 July insolation) from the AM China cave stalagmite δ 18 O record (Cheng et al., 2016). (d) Compariso of MD98-2165 δ 18 O sw-ivc and China cave stalagmite δ 18 O record (Cheng et al., 2016). (e) Comparison of MD98-2165 δ 18 O sw-ivc and precession parameter (Laskar et al., 2004).

The long, continuous δ 18 O sw-ivc of MD98-2165 makes it possible to examine the orbital time scale change of the interhemispheric thermal gradient. Spectral analysis reveals 100-ka, 41-ka and 23-ka components in the record of MD98-2165 δ 18 O sw-ivc . The 100-ka component from MD98-2165 δ 18 O sw-ivc is not related to the orbital eccentricity cycle. It is mainly caused by the increasing trend during MIS 6. As have been discussed in section 5.2, this long-term decreasing trend appears to reflect a weakening of ITF, allowing the saltier Indian Ocean water to increase its impact on δ 18 O sw-ivc at the site of core MD98-2165. There is both strong 41-ka and 23-ka components in MD98-2165 δ 18 O sw-ivc and China cave ∆δ 18 O, corresponding obliquity and precession bands. As shown in Figure 8, the 41-ka components filtered from the MD98-2165 δ 18 O sw-ivc and China cave ∆δ 18 O varied roughly in phase. However, these two 41-ka components both obviously lead the orbital obliquity cycle during the period of 0-90ka. This seems to suggest such 41-ka cycle in the MD98-2165 δ 18 O sw-ivc and China cave ∆δ 18 O is not correlated to the orbital obliquity. We cannot rule out possibility of such out phase relationship is from the bias by the millennial scale event during the spectral analysis. For example, the anomalously enrichment in MD98-2165 δ 18 O sw-ivc at ~16 ka and 60 is able to produce a 41-like cycle. As for at the precession band, there is a good correlation between the corresponding 23-ka components in MD98-2165 δ 18 O sw-ivc and China cave ∆δ 18 O and the orbital precession parameter. This good correlation could also be seen between the MD98-2165 δ 18 O sw-ivc and the orbital precession component (Figure 7e). The change of orbital obliquity and precession parameter can induce substantial difference in seasonal insolation received by the two hemispheres, thus affecting the seasonal interhemispheric thermal gradient as recorded by the cave δ 18 O records from the position near the seasonal boundary of ITCZ, such as Flores cave and China cave (Ayliffe et al., 2013;Cheng et al., 2016). However, the orbital obliquity and precession parameter change have negligible effects on the annual mean insolation difference received between the two hemisphere. As we considered MD98-2165 δ 18 O sw-ivc , Borneo cave δ 18 O and China cave ∆δ 18 O as an indicator of mean state of interhemispheric thermal gradient change, it means the corresponding 23-ka and 41-ka component is irrelevant to the direct insolation forcing.

Therefore, the AMOC is responsible for the hydroclimate change in the Indonesian part of IPWP on both millennia and orbital time scale.

Conclusions

We reconstructed a 270-ka long record of surface water δ 18 O corrected for the global ice volume effect (δ 18 O sw-ivc ) from sediment core MD98-2165 located in the tropical Eastern Indian Ocean. , 2002). (c) Comparison of δ 18 O sw-ivc obtained using sea level curve of (Waelbroeck et al., 2002) and by using MD98-2165 benthic-δ 18 O equivalent sea level curve. (d) Comparison of δ 18 O sw-ivc obtained by using (Waelbroeck et al., 2002) and the average of all results of δ 18 O sw-ivc in (b).

Chapter 5 : Indonesian Throughflow outflow variability over the last two climatic cycles : evidence from the eastern Indian Ocean

This chapter focuses on the variability of ITF transport over the last two climatic cycles. For this purpose, we analyze past changes of the thermocline water temperature in the exit of ITF into the Indian Ocean (derived from P. obliquiloculata Mg/Ca-TWT in core MD98-2165) and compare our results with various TWT records from the Indonesian Archipelago, including a recent record located in the main upwelling located to the West of Sumatra. These comparisons make it possible to better decipher the relative evolution of ITF strength in the context of the strong wind-driven upper ocean dynamics that affect most of the Indonesian Archipelago.

Indonesian Throughflow outflow variability over the last two climatic cycles: evidence from the eastern Indian Ocean

Xiaolei PANG 1,2 , Franck BASSINOT 1 , Sophie SEPULCRE 2 1 Laboratoire des Sciences Du Climat et de l'Environnement, LSCE/IPSL, CEA-CNRS-UVSQ, Université Paris-Saclay, Gif-sur-Yvette, France 2 GEOPS, Université Paris-Sud, CNRS, Université Paris-Saclay, Rue du Belvédère, 91405, Orsay, France.

In preparation

Abstract

The Indonesian Throughflow (ITF) is the main low-latitude return branch of the global thermohaline circulation. The heat carried by the ITF from the tropical western Pacific to the eastern tropical Indian Ocean strongly affects local and global climate changes. The magnitude and variability of its transport at different timescales are key factors in the inter-basin transfer of heat and salt and climate teleconnections between low and high latitudes. We present foraminiferal Mg/Ca-surface (SST) and thermocline water temperatures (TWT) covering the last 270,000 years in south of the Lombok Strait, one of the main exits of the ITF into the Indian Ocean. The P. obliquiloculata Mg/Ca-TWT changes reflect the balance between local winddriven upwelling and the ITF thermocline transport. The comparison of our records with those from other sites in the Indonesian areas make it possible to disentangle ITF changes from local variations associated to monsoon. Our results indicate that the ITF was weaker during MIS 8, MIS 6 and MIS 4-2, and was enhanced during MIS 7, MIS 5 and MIS 1. On the glacialinterglacial timescale, changes in TWT at our study site likely reflect sea level changes that modulate the local intensity of ITF through the modification of pathways and/or alter local SE monsoon winds by changing the land/ocean distribution. On orbital timescales, our results reveal the strong positive link between SE monsoon winds and ITF intensity.

Introduction

The Indonesian archipelago is a low-latitude oceanic pathway for the surface and subsurface water transport from the western tropical Pacific Ocean to the eastern tropical Indian Ocean.

The so-called Indonesian Throughflow (ITF) is the only low-latitude return branch of the global thermohaline circulation (Gordon, 2005;Talley, 2013). The three main exits of ITF into the Indian Ocean are Lombok Strait, Ombai Strait and Timor passages (Figure 1). After entering into the Indian Ocean, a sub-branch of the ITF travels towards Africa as part of the Indian South Equatorial Current (SEC) and can eventually exit the Indian Ocean to the Atlantic Ocean through the Agulhas leakage around the southern tip of Africa (Beal et al., 2011;Gordon, 2005;Peeters et al., 2004). Another branch of the ITF joins the Leeuwin Current (LC), which travels southward along the west coast of Australia, and reaches as far south as Tasmania (De Deckker et al., 2012;Gordon, 2005). Thus, the ITF is a key contributor to a circulation system that transports oceanic heat from the Pacific Ocean to the Indian Ocean and from low latitudes to the high, southern latitudes of the Indian Ocean and, ultimately, to the Atlantic Ocean (through the Agulhas leakage). Despite the key role played by ITF as a heat and fresh water carrier, its variability and its exact role in driving regional and global climate changes at millennial and orbital timescales are still poorly known due to the difficulty of constraining the strength of the ITF in the past, and the paucity of sedimentary records allowing to reconstruct hydrographic properties of water masses entering the Indian Ocean. These difficulties traduce the extreme complexity of the ITF structure and time-evolution, which are difficult to properly grasp even today. Datasets are from WOA13 (Locarnini et al., 2013).

Although the total ITF flux extends down to about 700 m in the water column, the main core of modern ITF transport is located at the thermocline depth (Gordon, 2005;J Sprintall et al., 2009). Modern observations averaged other several years indicate that ~15 Sv (1 Sv = 10 6 m/s) of water are carried by the ITF to the Indian Ocean (Sprintall et al., 2009). But this averaged number masks a large amplitude variability (up to ± 1.5 Sv compared to the average 15 Sv flux) On an inter-annual timescale, the hydrographic time series available for reconstructing ITF shows a considerable complexity when comparing the upper layer flux (0-150 m) at the three main exit pathways (Sprintall & Révelard, 2014). When removing the long-term trend and seasonal signals, hydrographic datasets reveal that the expected response of a reduced ITF to the relaxation in the Pacific wind field during El Niño events only occur for the Ombai Strait (Figure 3; upper left). But the ITF clearly collapses (negative anomalies) when El Niño events are coeval with negative Indian Ocean Dipole (IOD) events (Figure 3; upper right). In the case of solo La Niña events, only the Timor Passage shows a positive transport anomaly (Figure 3; lower left), whereas the co-occurrence of La Niña events and negative IOD results in positive transport anomalies in the three main pathways, and particularly in the Timor Sea (Figure 3; lower right). Thus, clearly, the ITF flux depends upon the intensity of zonal circulation but in a complex way, involving the inter-connection of both the Pacific (ENSO) and Indian Ocean (IOD) systems.

There are no direct ways of estimating past changes in the flux of ITF. Thus, the few existing paleo-reconstructions of ITF obtained from marine sediment cores were based on the reconstruction of surface and sub-surface temperatures as indicators of paleo-ITF dynamics (Fan et al., 2018;Holbourn et al., 2011). For instance, Holbourn et al. (2011) have estimated past changes in ITF dynamics over the last glacial-interglacial cycle by reconstructing the thermocline structure of the upper water column using a marine sediment core located off the coast of NW Australia, in the SW Timor Sea. Their approach is based on the comparison of (i) Mg/Ca-derived surface temperatures obtained from the planktonic shallow dwelling species Globigerinoides ruber, with (ii) Mg/Ca-derived thermocline temperature obtained from the deeper dwelling species Pulleniatina obliquiloculata. The problem with such an approach, is that both seasonal and inter-annual temperature patterns associated to ITF are partially masked by the superimposition of local changes in temperature related to the development of seasonal upwelling cells forced by monsoon-winds.

Along the southern coast of Lesser Sunda, Java and the adjacent island of Sumatra, north west of the Timor Sea, a major seasonal upwelling develops (Steinke et al., 2014;Susanto et al., 2001). During the southeast (SE) monsoon season (June to September), southeasterly winds blowing along the coasts generate an upwelling through Ekman pumping, inducing a cooling of the thermocline water temperature (TWT) (Figure 2d). Conditions reverse during the northwest (NW) monsoon season (December to February), when the regional atmospheric Using a Mg/Ca-TWT record from the sub-surface dwelling planktonic foraminifer Pulleniatina obliquiloculata picked along sediment core SO139-74KL (Figure 1), Wang et al. (2018) showed that TWT variations at millennial and orbital timescales off the coast of western Sumatra had always been chiefly controlled by the upwelling intensity over the last 300 ka.

Interestingly, changes in the TWT of core SO139-74KL exhibit significantly different features compared to TWT records from non-upwelling regions in the Indonesian archipelago, which are thought to be affected by variations of the ITF (Wang et al., 2018). These results suggest that comparing surface and thermocline records off the west coast of Sumatra with records from the main exit flow of ITF may actually help us to disentangle to some extent the ITF variability from the seasonal, monsoon driven variability. Of course, the complex inter-connection of zonal and meridional atmospheric circulation is such that it is illusive to fully extract pure ITF from pure monsoon-related signals. Nevertheless, we believe that combining the upwellingdominated TWT record of Sumatra (SO139-74KL) with records from more complex areas, showing a mix influence of ITF and monsoon-driven variability, can help deciphering the periods of major ITF changes.

For such purpose, the eastern side of the Java-Sumatra upwelling zone, west of the Ombai Passage, seems a good potential area to compare with paleo-records located off the west coast of Sumatra (SO139-74KL; Wang et al., 2018) and within the Timor Sea (MD01-2378; Holbourn et al., 2011). This "transitional area" located west of the Ombai Passage (and SE from the Lombok Strait) is under the mixed influence of seasonal upwelling (Figure 2) and ITF transport (Figure 1). Thus, it is our contention that a long paleoceanographic record located there should make it possible, by comparison to the SO139-74KL record off Sumatra, to disentangle millennial to orbital changes associated to the lateral advection of warm ITF water, from those resulting from wind-driven, vertical mixing of sub-surface waters.

In this study, we address these issues by looking at SST and TWT changes from core MD98-2165 (Figure 1) covering the last 270 ka. Temperatures were obtained by analyzing Mg/Caratios on the surface dwelling planktonic foraminifer G. ruber and the thermocline dwelling species P. obliquiloculata. In this chapter, we discuss the changes in the ITF transport and its influence on thermocline structure and temperature.

Modern oceanographic settings in the studied area

The dynamics of ITF and monsoon-driven upwelling at seasonal and inter-annual timescale have direct impact on the modern TWT variations recorded (i) west of the Ombai Striat (at the location of Core MD98-2165), in the south Java upwelling system and (ii) off the coast of central Sumatra (at the location of Core SO139-74KL), in the heart of the SE monsoon upwelling system, but more protected from the main ITF flow Figure 1). Modern SST variations, west of the Ombai passage, south of Java, show regular seasonal changes (Figure 5, red line) and range between 25 to 31°C, with an average of 28.2°C (average of temperatures recorded over 30 years, from 1985 to 2015; data from the Simple Ocean Data Assimilation (SODA) (Carton & Giese, 2008). By contrast, SST variations of the coast of central Sumatra, display smaller seasonal changes, around a slightly higher mean annual SST of 28.6 °C (Figure 5). The smaller decrease of SST along Sumatra compared to the south Java area during the austral winter upwelling season results from a "barrier layer" off Sumatra. This barrier layer separates the base of the mixed layer from the top of thermocline, preventing the upwelling of cold water into the surface (Qu & Meyers, 2005). TWT are lower during the austral winter season, when the wind-driven (SE monsoon), seasonal upwelling is strong. Modern TWT vary synchronously in both areas, clearly showing the strong influence of the upwelling dynamics on a seasonal timescale. West of the Ombai passage (and SE of Lombok Strait), the area is also affected by the lateral ITF thermocline transport of warm water (Sprintall et al., 2009). This effect may explain the somewhat lower amplitude of TWT changes, the TWT temperatures being somewhat modulated and forced to remain more strictly around a temperature characteristic of the ITF thermocline flow.

Reconstructing the past thermal structure of the upper ocean

As seen above, reconstructing surface and sub-surface temperature changes in the past are key elements if one wants to reconstruct the relative variability of ITF and upwelling dynamics at timescales exceeding those offered by instrumental records, and in particular to decipher millennial to orbital changes. The difference in Mg/Ca-temperature between the mixed-layer species Globigerinoides ruber and thermocline species P. obliquiloculata were used by several authors to track changes in the upper water column stratification (Dang et al., 2012;Hollstein et al., 2018;Mohtadi, Oppo, Lückge, et al., 2011;Jian Xu et al., 2010). Sediment trap experiments in the seasonal upwelling area off south Java indicate habitats of 0 -30 m for G. ruber and 60 -90 m for P. obliquiloculata (Mohtadi et al., 2009). The study of surface sediment in the Lombok basin, near core MD98-2165, indicates a slightly deeper 100 m habitat for P. obliquiloculata. Although higher fluxes of foraminifer are expected during the upwelling season, the flux-weighted shell geochemistry data (e.g., oxygen isotopes δ 18 O and Mg/Ca) are consistent with average values in surface sediment samples (Mohtadi et al., 2009). G.ruber shell geochemistry in core tops represents annual mean surface conditions (Mohtadi et al., 2009;Mohtadi, Oppo, Lückge, et al., 2011). P. obliquiloculata sediment shell geochemistry, however, is biased by the higher fluxes during the austral winter season, and tends to record austral winter (the upwelling season) upper thermocline conditions (Mohtadi, Oppo, Lückge, et al., 2011).

In the present study, we will present G. ruber SST and P.obliquiloculata TWT records from core MD98-2165 and we will compare them with similar records obtained off the coast of central Sumatra (SO139-74KL) and in the center of the Timor Sea (MD01-2378) in order to better constrain past variability of ITF and its potential impact on the hydrology of the Indian Ocean.

Material and methods

Sediment core MD98-2165 and age model

Sediment core MD98-2165 (9.65S, 118.34E; water depth of 2100m) was retrieved from the south east of the Lombok Strait and west of the Ombai Passage, two main exits of ITF in the eastern tropical Indian Ocean. The core was retrieved during the Images IV cruise of the French R/V Marion Dufresne (Figure . 1). The core is 42.3m long. Only the upper 34.5m are studied in this paper. An initial age model had been constructed by using absolute radiocarbon ages (on the planktonic foraminifer G. ruber) and benthic δ 18 O foraminiferal stratigraphy (Levi et al., 2003;Levi et al., 2007;Waelbroeck et al., 2006). In this initial age model, calendar ages were computed by using the 1998 marine calibration curve with a surface reservoir age of 400 years (Waelbroeck et al., 2006). Here, we used the latest MARINE13 age calibration curve (Reimer et al., 2013) to convert AMS 14 C dates to calendar age through Calib 7.10 software (Table 1).

Marine reservoir ages were set by default to 405 years. The new generated calendar ages are consistent with previously published ages for 14 C dates younger than 16 ka, and ~200-500 years older for 14 C age in the range 16-26 ka. For the older part of the core, we established a new age model by aligning the benthic δ 18 O record of MD98-2165 to a global benthic δ 18 O stack. We used the LR04 stack (Lisiecki & Raymo, 2005) in the interval 270-150 ka, and the revised LS16 stack (Lisiecki and Stern, 2016) in the interval 150-0 ka. Over the part of the MD98-2165 record where age model rests upon this tuning to reference isotopic records, we used 5 tie points.

Depths were converted to ages by simple linear interpolations between bracketing AMS 14 C calendar dates or δ 18 O tie points. Based on this age model, the 3450 cm record of core MD98-Table 1. AMS 14 C dates and corresponding calendar ages of core MD98-2165 .* 14 C dates were taken from Waelbroeck et al. (2006).

Mg/Ca-temperature reconstructions in core MD98-2165

In this study, G. ruber Mg/Ca derived SST datasets are taken from previous studies (0-25 ka, Levi et al., 2007;25-270ka, Chapter-4 paper of this thesis). Two anomalous higher SST data at core top interval of 60-70 cm in the study of Levi et al. (2007) are excluded. In this paper, we provide new P. obliquiloculata Mg/Ca data from the interval of 110 -3450 cm corresponding to ~1.5 -270ka. Mg/Ca data for P. obliquiloculata were obtained as part of this work. For measuring P. obliquiloculata Mg/Ca ratios, Core MD98-2165 was sampled at 4-12 cm interval in the 110-3450 cm depth range and totally 410 samples were analyzed. Samples wet sieved and dried overnight and about 30 tests of P. obliquiloculata were picked from the 250-355 µm size fraction. Foraminiferal tests were gently crushed between two glass plates under a binocular.

Crushed tests were then transferred to acid-cleaned 500 µl centrifuge tubes for further cleaning.

Samples were cleaned following the method described in Barker et al., (2003). After cleaning, each sample was dissolved within 0.075M HNO 3 and then centrifuged for removing undissolved particulate impurities. The supernatant solution was transferred to a new acidcleaned 500 µl centrifuge tube. Mg/Ca ratios were measured using an inductively coupled Thus, no Mg/Ca data was rejected from our dataset.

Conversion of Mg/Ca ratios to temperatures and consistency of datasets

Mg/Ca-based seawater temperatures were compared between different sediment cores in this study. As shown in Table 2, the Mg/Ca from the different studies were measured from slightly different foraminifera shell sizes and by using different cleaning methods. Xu et al. (2010) demonstrated that for G. ruber and P. obliquiloculata species from 250-315 µm and >315 µm size fractions, the Mg/Ca ratios did not exhibit significant differences. Thus, we do not take into account any sample size effect on the Mg/Ca ratios. However, the use of different cleaning methods for Mg/Ca analysis, specifically with or without the reductive step, lead to different Mg/Ca-based temperature estimates (Barker et al., 2003;Rosenthal et al., 2004

Results

Decoupled SST and TWT signals at site of core MD98-2165

Mg/Ca ratios of P. obliquiloculata in core MD98-2165 vary between 1.6 and 3.1 mmol/mol, corresponding to TWT changes between 17.8 and 25.1 °C (Figure 6). The core top TWT (at ~1.5ka) is ~ 20.4°C which is near to the anomalously lowest modern temperature of ~21°C at the water depth of 75m. The SST from the same core top interval is ~27.8°C, similar to the annual mean surface temperature ~28.2°C. Consequently, the difference ΔT (SST-TWT) is ~7.4°C (at ~1.5ka). (Levi, 2003;Levi et al., 2007;Waelbroeck et al., 2006) and global benthic δ 18 O stack (black line) (Lisiecki & Raymo, 2005;Lisiecki & Stern, 2016).

On longer timescales, TWT displays similar glacial-interglacial variations as SST, with relatively lower values during the colder marine isotope stages (MIS) 8, MIS 6 and MIS2-4 than the warmer stages MIS 7, MIS 5 and MIS 1. Both SST and TWT show a long-term decreasing trend from ~130 to 20ka. Although there is an overall good correlation on glacialinterglacial timescale, TWT show substantial amplitude change differences compared to SST.

This can be readily seen from the Mg/Ca-ΔT records. Mg/Ca-ΔT ranges between 3 and 7°C,

showing generally lower value during the period 270 -130 ka than the following period 120 -30 ka and the late Holocene. Mg/Ca-ΔT displays particularly lower peak values glacial terminations and deglaciations.

In detail, the TWT is ~21°C at ~270ka which is warmer than the core top value of 20.4°C (~1.5ka). In contrast, SST indicates ~2°C colder at ~270ka (~25.6°C) than ~1.5ka (~27.8°C).

As a result, the corresponding ΔT is ~4.5°C which is larger than core top (~1.5ka) ΔT(~6.8°C), Similarly, our data show that there is no strong covariance either between TWT and regional surface temperature changes at the glacial/interglacial timescale. The somewhat decoupled SST and TWT variability likely indicates the persistence, at the G/IG timescale, of a barrier layer which did not make it possible a strong vertical coupling, neither through a downward flux of heat from the surface water (i.e. mixed layer) nor through the upwelling of cold water to the surface. This suggests that long term TWT at the site of Core MD98-2165 chiefly reflect the impact of upwelling intensity and ITF (flux and/or temperature) on the thermocline.

Spectral analysis: TWT changes lead SST changes at orbital time scales

Spectral analyses performed on TWT and SST records of MD98-2165 reveal 100-ka, 41-ka and 23-ka periodic components (Figure 7). Despite showing clearly different patterns over long time scale, the spectral power of both records is dominated by a strong ~ 100-ka cycle. In order to enhance the 41-ka and 23-ka periodic components, lower frequency components (periods longer than 50-ka) were removed from SST and TWT records by notched filtering (Gaussian) at a central frequency of 0.01 and bandwidth of 0.01 with using the software Analyseries software (Paillard et al., 1996). After removing the low variability pattern, our data reveal a staggering co-variance between higher frequency residuals (Figure 7d). Both TWT and SST residual records are characterized by comparable 41-ka obliquity and 23-ka precession periodic components (Figure 7e). The residual of TWT displays a more intense component at the 23-ka (precession) cycle than at the 41-ka cycle characteristic of obliquity orbital cycle (Figure 7f).

Moreover, the TWT record shows a significant half-precession cycle (11-ka) that is absent in the SST record (Figure 7). The 11-ka half-precession cycle is a signal that has been exclusively observed in equatorial (and, to a lesser extent, the intertropical) paleoclimate records driven by insolation (Ashkenazy & Gildor, 2008;Berger et al., 2006). This suggests that TWT variations at the site of core MD98-2165 convey a direct imprint of tropical insolation changes. at a central frequency of 0.01 and bandwidth of 0.01 performed using the AnalySeries software (Paillard et al., 1996); (e) spectral analysis results for the notched SST (<50 ka); (F) spectral analysis results for the notched TWT (<50 ka). Spectral analysis was performed by the method REDFIT (Schulz & Mudelsee, 2002) through software of PAST (Harper & Ryan, 2001); a Welch type window was used with setting of oversample = 1 and segments = 1; red-noise spectrum (black line) and 90% confidence levels (pink line) were showed in (b), (c), (e) and (f).

The main orbital timescale components were filtered out from the original records of SST and TWT in order to readily observe how these components vary through time in both records (Figure 8). At eccentricity, obliquity and precession periodic cycles, the TWT variations consistently lead SST change. As filtering results are somewhat biased at the upper and lower limits of a given series, we cannot fully rely on the uppermost and lowermost offsets observed between the two datasets. But in the middle part of the records, this lag of TWT over SST is real, suggesting that, at orbital time scales, climate processes that control TWT react to insolation changes earlier than the climate processes that force SST changes in the area of Core pass filtering was performed by using software of AnalySeries (Paillard et al., 1996). The corresponding eccentricity component was filtered out from SST and TWT at frequency 0.008±0.005 ka -1 ; the corresponding obliquity component was filtered out from SST and TWT at frequency 0.026±0.004 ka -1 ; The corresponding precession component were filtered out from SST and TWT at frequency 0.044±0.004 ka -1 .

Comparison of MD98-2165 and SO139-74KL SST and TWT records

As shown in Figure 9, SST at site MD98-2165 and SO139-74KL show high similarity on orbital to glacial-interglacial timescales (Figure 9a). However, SST is generally lower at site of core SO139-74KL than at site of core MD98-2165 all over the past 270ka. This is particularly noticeable during MIS 7 (at ~240ka, 215ka) and MIS 5 (at ~130ka and 80ka), when SST at site of core SO139-74KL was more than 1°C lower than at site of core MD98-2165. The lower temperature at site SO139-74KL is also evident in heavier G. ruber δ 18 O of SO139-74KL (Figure 9d).

P. obliquiloculata Mg/Ca-derived TWT show more significant difference between cores SO139-74KL and MD98-2165 (Figure 9b). Generally, during MIS 8, MIS 7, MIS 5 and MIS1, TWT at the site of core MD98-2165 was at least 1°C and up to 6°C higher than at the site of core SO139-74KL (Figure 9d). By contrast, the TWT gradient between these two core sites was relatively small during MIS 6 and MIS 2-4. Overall, the TWT gradient between the two sites reveals a very and strong variability at interglacial and orbital timescales (Figure 9d). TWT in core MD98-2165 also show remarkable differences compared to TWT from core SO139-74KL during warm stages MIS 7, MIS 5 and MIS 1. These long-term differences in the two TWT signals contrast with the strong seasonal and inter-annual co-variance of TWT during modern conditions (Figure 5). At present, at both sites, it has been estimated that P.

obliquiloculata has an average calcification depth around 75 m (Mohtadi, Oppo, Lückge, et al., 2011). The study of core top sediments suggests that Mg/Ca in P. obliquiloculata represents austral winter season conditions (Mohtadi, Oppo, Lückge, et al., 2011) due to the increased nutrient input associated with the seasonal upwelling, which results in a higher production flux of P. obliquiloculata during the winter season compared to the rest of the year (Mohtadi et al., 2009). Thus, during the past, one can expect that changes in the upwelling strength are also strongly imprinted in P. obliquiloculata Mg/Ca-derived TWT in the study area. Indeed, Wang et al. (2018) have concluded that the TWT of core SO139-74KL reflect past changes in the upwelling intensity off the west coast of Sumatra, with colder TWT reflecting strengthened local upwelling.

At present, the upwelling along the south coast off Java-Sumatra is induced by the southeast (SE) monsoon wind, which drives offshore Ekman transport during the austral winter season (i.e. boreal summer season) (Susanto et al., 2001). Hydrographic data and satellite observations of chlorophyll production reveal that the upwelling along Java and Sumatra propagation spatially during the course of austral winter, the core of the upwelling migrating slowly from the South of Java to the West of central Sumatra in response to complex mechanisms resulting in part from the slight evolution of intensity and direction of SE winds through time (Susanto et al., 2001). On the long time scale available from marine sediment cores, the SE winds have always been a permanent feature of seasonal wind circulation in this region. It is the authors' contention that this remanent wind system has always induced strongly coupled offshore Ekman transport at the sites of core SO139-74KL and core MD98-2165. insolation difference between 20N and 20S during the period of June to August (Laskar et al., 2004). (c)Total organic carbon (TOC) record from core SO139-74KL (Lückge et al., 2009), and JJA-Insolation. (d) relatively sea level change (Waelbroeck et al., 2002).

The seasonal reversal change in monsoon wind regime is ultimately linked to the latitudinal and seasonal evolution of insolation radiative between the north and south hemispheres, driven by parameters of Earth's orbit around the Sun. On the orbital scale, Lückge et al. (2009) showed that the intensity of the costal upwelling driven by SE monsoon wind is correlated to the Northern hemisphere summer insolation, with enhanced upwelling associated to periods of increased boreal summer insolation. In accordance to the seminal approach of (Rossignol-Strick, 1985), we prefer to consider the global monsoon wind system as driven by the thermal gradient between the northern and southern hemispheres. When only taken into account external forcing, such a gradient can be simply approached by assuming that is chiefly reflects direct contrast in insolation between latitudes and seasons. We calculated the mean insolation difference between 20°N and 20°S during boreal summer season (June to August, JJA-Insolation hereafter)

as an indicator of temperature differences between the tropical regions of the two hemispheres induced by orbital changes (Figure 10b). Increased JJA insolation (boreal summer, austral winter) will favor enhanced SE monsoon winds which is evident by the good correlation with total organic carbon record (TOC) of core SO139-74KL (Lückge et al., 2009) (Figure 10c).

Interestingly, there are striking positive relationships between ΔTWT -the TWT difference between core MD98-2165-TWT and core SO139-74K-TWT -and JJA-Insolation (Figure 10b), with increased thermocline temperature gradient between these two sites corresponding to increased JJA-Insolation. This result suggests, at this orbital time scale, increased SE monsoon driven Ekman offshore surface transport bring cold water to the thermocline depth at site SO139-74KL, but relatively warmer water to the site MD98-2165. This can be readily seen in Figure 11, with peak JJA-Insolation corresponding to peak TWT in MD98-2165. We believe that the different response of P. obliquiloculata Mg/Ca derived TWT to orbital forcing at the sites of core SO139-74KL and core MD98-2165, which results in a clear orbitally-forced pattern in the TWT gradient, is an indication that another factor than upwelling also impact the TWT at the site of core MD98-2165. Our prime suspect for such additional control on TWT, is past changes in ITF thermocline flux and/or temperature. In order to better separate the various, we use period coving the last glacial maximum to the upper Holocene and compare the TWT records of cores MD98-2165, SO129-74KL and MD01-2378 (in the Timor Sea; Figure 12) with reference also to modern conditions. Over the last 25 ka interval, the chronology rests upon radiocarbon age tie-points, making it easier and more robust the comparison between the three records and with the insolation forcing. (Mohtadi, Oppo, Steinke, et al., 2011)).

At present, the ITF at all three exit straits (Lombok strait, Ombai strait and Timor passage)

shows maximum transport in the thermocline (~100 -150m) during the SE monsoon season (Sprintall et al., 2009). Because SE monsoon winds drive strong Ekman divergence in the eastern Indian Ocean side, it also helps maintaining a relatively lower sea level than in the western Pacific Ocean, therefore promoting the ITF transport towards the Indian Ocean (Sprintall et al., 2009). Core SO139-74KL is not located in the main ITF stream, whereas core MD98-2165 and MD01-2378 are (Figure 1). The SE monsoon driven offshore Ekman transport at site MD98-2165 is compensated by both vertical cold deep upwelling water and horizontal advection of ITF thermocline water. Although ITF flux is maximum during the SE monsoon, the thermocline water temperature in the vicinity of MD98-2165 (taken at 75 m in Figure 2d) appears chiefly controlled by the upwelling of cold deep water. Core-top sediments study also shows that P. obliquiloculata Mg/Ca is comparable between site SO139-74KL and site MD98-2165 (Mohtadi, Oppo, Lückge, et al., 2011), further suggesting that ITF have not exerted the main control on P. obliquiloculata shell geochemistry over the recent past. These results suggest that under modern conditions, the ITF transport (flux and/or heat) is not able to strongly imprint its signature over that of the upwelling system at site of core MD98-2165.

By contrast, over the early Holocene (~10ka; Figure 12), the SE monsoon strength and the wind-driven upwelling was at peak and much stronger than today as shown by the percentage of Globigerina bulloides in core GeoB10053-7 (Mohtadi et al., 2011). The stronger upwelling and without any strong upwelling, P. obliquiloculata Mg/Ca is supposed to record annual mean conditions at site MD01-2378 (Holbourn et al., 2011;Zuraida et al., 2009). P. obliquiloculata Mg/Ca derived TWT of MD98-2165 and MD01-2378 are coeval in the 16 to ~ 4 ka interval, with the same peak centered around 10 ka (Figure 12). The fact that TWT of MD98-2165 and MD01-2378 are so similar during this time interval suggests that they either record the same season, or that the slightly higher flux of P. obliquiloculata during the SE season at MD98-2165 did not significantly overshadowed the ITF temperature signal. This would suggest also that the flux of ITF may have been increased to compensate for the increased upwelling activity in the area. A question remains: was such an increase in ITF a "seasonal" enhancement linked to the stronger SE monsoon prevailing during the early Holocene, or was such an increase associated to a more "La Niña like" situation, with more intense and more frequent periods associating La Niña zonal intensification and negative IOD events?

Taken at face value, our data and the comparison with TWT from MD01-2378 and SO139-74KL suggest that during the early Holocene, around the peak in northern latitude summer insolation, when SE monsoon winds and upwelling intensity reached a peak, the ITF thermocline water was warmer and the its flux was increased. It is our contention that the same conclusion can be also applied to explain the ΔTWT peak oscillations change MIS 7 and MIS 5. We suggest that increased SE monsoon winds by producing enhanced upwelling at core site SO139-74KL and promoting warm ITF transport at core site MD98-2165 is responsible for the change of ΔTWT.

2 Sea level modulation of ITF transport

The ΔTWT as well as the ΔTWT amplitude variation are generally small during the cold stages MIS 8, MIS 6 and MIS 2-4 (Figure 10). These change in ΔTWT do not follow the JJA-insolation changes in amplitude. For example, during the MIS 7 to MIS 6 transition period (185 -170ka), JJA-Insolation increase ~ 40 w/m 2 , which is comparable to change during MIS 6 to MIS 5 transition period (135 -125ka), however the ΔTWT only changed by ~2°C compared to ~5°C

ΔTWT change during MIS 6 to MIS 5. Smaller ΔTWT correspond to relatively lower sea level periods. Sea level change can exert effect on ΔTWT through two ways. First, lower sea level can narrow the ITF pathways and exits of straits, therefore probably decrease the ITF transport at site MD98-2165 and thus depress the ΔTWT. Second, lower sea level can expose shallow shelf around the Indonesian seas region, changing land-sea thermal gradient and regional albedo, thus influencing the SE monsoon strength as well as the zonal atmospheric circulation (DiNezio et al., 2018). The total organic carbon (TOC) record, a proxy of local upwelling strength (Lückge et al., 2009), at site SO139-74KL indicate that the amplitude of the upwelling was only (Wang et al., 2018) and MD06-3067 (Bolliet et al., 2011).

Summary

P. obliquiloculata Mg/Ca-derived TWT at site MD98-2165 (South of Java, west of the Ombai Passage) reflect the balance between ITF transport and local upwelling. Mg/Ca-derived temperature records covering the last 25 ka show that TWT at the site of core MD98-2165 varied continuously between two end-members: (i) TWT from the Timor Sea (core MD01-2378, within the main ITF pathway) and (ii) TWT from a pure upwelling record, off the west coast of Central Sumatra (core SO139-74KL). From 25 to about 16 ka, TWT at the site of core MD98-2165 were similar to TWT from core SO189-74KL located in the Sumatra upwelling, away from the ITF main pathways. This indicates that the upwelling signal dominated TWT at these two sites, suggesting a reduced impact of thermocline ITF. In the interval ~ 16-4 ka, and more strikingly at the peak of boreal summer insolation (austral winter), around 10 ka, the TWT at site of core MD98-2165 were much warmer that at the site of core SO139-74KL and similar

to TWT from the central Timor Sea (MD01-2378), suggesting a dominant ITF effect (i.e.

increased ITF flux and/or temperature). Interestingly, the ITF signal overshadowed the upwelling effect in core MD98-2165, although upwelling index from cores located off the Sumatra margin indicate that the Java/Sumatra upwelling reached its maximum intensity over this period. From 4 ka upward, TWT from at the site of core MD98-2165 tends to become progressively similar to those from the upwelling site SO139-74KL.

Our data either suggest that (i) the strength of monsoon SE winds, paced by insolation, was the main driver of ITF intensity over the last 25 ka (i.e. similar to the modern seasonal pattern of ITF variability) and/or (ii) that ITF changes were ultimately driven by variations in the zonal atmospheric circulation (i.e. similar to the mechanisms that control, today, the anomalies of ITF at the inter-annual timescale, with decreased ITF flux during El Niño/IOD+ events and increased ITF flux during La Niña/IOD-events). If this second hypothesis prevails, and if modern observations (i.e. Sprintall and Révellard, 2014) can be applied to longer timescales, our data would tend to suggest that the early Holocene was characterized by more frequent and/or more intense episodes combining La Niña and IOD-events. However, the tight links that exist between the meridional (monsoon related) and the zonal atmospheric circulation patterns, together with our incapacity to resolve short periods (plural-annual to seasonal) in marine sediment records, make it is impossible to properly decipher which of our two hypothesis prevail and which mechanisms ultimately control ITF intensity.

As seen above, difference in TWT between the site of core MD98-2165 (South of Java, two (2) We showed that weak correlations and sometime nearly anti-phased relationships exist between core MD98-2165 δ 18 O sw-ivc and the nearby Flores cave stalagmite δ 18 O record, clearly indicating that past salinity changes in the Timor Sea were not chiefly controlled by changes in the local precipitation regime. Instead, we showed that core MD98-2165 δ 18 O sw-ivc varied coincidently with Borneo cave stalagmite δ 18 O record, and both records are strongly correlated also with millennial-scale changes in a China cave stalagmite record that had been corrected for major orbital oscillations (∆δ 18 O). Our results underline, therefore, (i) the existence of coherent modes of hydroclimate changes recorded over most of the Australian-Asian monsoon region, and (ii) the fact that the surface water exiting the Indonesian Archipelago as part of ITF carries a δ 18 O sw signal which is coherent with northern hemisphere records (i.e. Borneo, China)

and sometimes anti-phased with local, southern hemisphere precipitation changes (i.e. Flores). higher during the early Holocene (~10 ka) than at most recent past in our records at ~1.5 ka.

This higher early Holocene Mg/Ca-TWT was attributed to the increased temperature and/or flux of ITF water in conjunction with the weakened upwelling over MD98-2165 (in a period during which upwelling intensity was higher at west of Sumatra). Whether the upwelling was Chapter 6 : General conclusions and perspectives 168 weaker at the site of MD98-2165 or whether it was strong but its effects were overshadowed by the higher ITF temperature and/or flux is still uncertain and could not be properly solved with our data. The scenario of a weak local upwelling during a period of strong local SE monsoon winds can be tested by studying the total organic carbon (TOC) in the same core MD98-2165. Because, even though the much warmer ITF thermocline water during the early Holocene (~10ka) can cancel the cold upwelling water effect in temperature, the ITF is not able to cancel the increased nutrient input along with the upwelling water. Thus, the TOC record can provide evidence of the upwelling change, and thus provide support to assert the ITF flux change.

Perspectives: ITF effect on Indian Ocean and low-and high-latitude connections

The ice-volume change corrected local seawater δ 18 O sw-ivc and G. ruber Mg/Ca-SST from core MD98-2165 represent surface water temperature and salinity change of ITF surface water. The study of P. obliquiloculata Mg/Ca-TWT provides constraints about the ITF thermocline water temperature and as well as the ITF flux change. Together, these results recorded the ITF water variations in both heat and salt that was injected to the Indian Ocean. At present, ITF thermocline transport -averaged over the whole water depth where it has an effect -tends to cool the Indian Ocean subsurface water mass, due to relatively lower ITF thermocline water temperature [START_REF] Gordon | Cool Indonesian throughflow as a consequence of restricted surface layer flow[END_REF]. Could this had been different in the past, when ITF thermocline was much warmer than today? Did the ITF warm or cool the Indian Ocean? As enhanced SE monsoon winds (and/or changes in zonal atmospheric circulation) modulate the ITF flux, then is it conceivable that increased ITF heat transport to the Indian Ocean may have exerted negative feedback on the local SE monsoon? To answer these questions, one needs to trace the trajectory of ITF in the central Indian Ocean and reconstruct the upper water thermal structure along the pathway of ITF in Indian Ocean.

As the upper branch of global thermohaline circulation (Gordon, 2005;Talley, 2013), the ITF plays critical role in the global climate change by connecting the low-and high-latitude through oceanic circulation. The ITF water eventually leaves Indian Ocean through the Agulhas Currents at the southern tip of Africa (Beal et al., 2011), and then joins into the Atlantic Meridional Overturning Circulation (AMOC) to complete the loop of the global thermohaline circulation (Talley, 2013). Several questions can be raised here based on the results of this PhD work: (i) could the hydroclimate change in the IPWP (revealed by the δ 18 O sw-ivc record of MD98-2165), have exerted significant effects on the southern, high-latitude Indian Ocean, and have significantly contributed to modify the spilling out of water in the Agulhas leakage system?

(ii) could the heat transported by the ITF from the IPWP to the south high-latitude have impacted the temperature changes at high latitudes and therefore influenced the latitudinal thermal gradients?

Preliminary efforts have been made to answer these questions by comparing local seawater δ 18 O sw-ivc record between the IPWP region and Agulhas Currents region. As can been readily

seen in Figure 6.3, local seawater δ 18 O sw-ivc record varied coherently between the IPWP region (core MD98-2165) and the Agulhas currents region (core MD02-2594) on both millennial to orbital timescales. As ITF is the source water for the Agulhas leakage (Beal et al., 2011), the simple explanation is that the salt transported by ITF from IPWP region control the salt leakage from Indian Ocean to the Atlantic Ocean. However, one another possibility cannot be ruled out is that similar regional precipitation at the two regions were the result of atmospheric circulation changes forced by the interhemispheric thermal gradient.

  climatiques passés. Une telle connaissance est cruciale pour fournir de meilleures contraintes sur le changement climatique et anticiper son évolution future sous un forçage radiatif différent. Dans le présent travail, j'aborderai plus particulièrement le rôle et les changements des mécanismes des basses latitudes associés au cycle hydrologique et au transfert de chaleur et de sel par la grande boucle de circulation thermo-haline. Avec l'augmentation du nombre d'observations paléoclimatiques et la meilleure compréhension du climat actuel, une image globale a émergé, qui souligne l'importance des interconnexions atmosphériques et océaniques. La redistribution de la chaleur et de l'humidité par les régimes de circulation atmosphérique et océanique est considérée comme un mécanisme clé pour comprendre et anticiper l'évolution du climat mondial et ses impacts régionaux ou locaux. Cette redistribution est affectée par de nombreuses boucles de rétroaction complexes et des interactions à distance.

  The sea surface temperature (SST) and thermocline water temperature (TWT) were reconstructed by the foraminiferal Mg/Ca paleothermometry. The sea surface salinity (SSS) was evaluated by reconstructing δ 18 O of seawater through the combination of Mg/Ca-T° and δ 18 O of planktonic foraminiferal tests. Before using Mg/Ca paleothermometry, the impact of cleaning methods (with or without a reductive step to remove oxides) were evaluated on planktonic foraminiferal Mg/Ca analyses. This work is mandatory if one wants to combine accurately Mg/Ca-based temperatures obtained from different studies, which used different cleaning methods in the IPWP region.

Figure 1 . 1

 11 Figure 1.1 The region of Indo-Pacific Warm Pool. It is defied by the annual mean temperature of 28 °C isotherm line.

Figure 1 . 2 .

 12 Figure 1.2. Upper panel: annual mean surface temperature; middle panel: monthly mean surface temperature of January; lower panel: monthly mean surface temperature of July. Dataset from WOA13 (Locarnini et al., 2013).

Figure 1 . 3 .

 13 Figure 1.3. Upper panel: annual mean temperature at water depth 100m; middle panel: monthly mean temperature (100 m) of January; lower panel: monthly mean temperature (100 m) of July. Dataset from WOA13 (Locarnini et al., 2013).

Figure 1 . 4 .

 14 Figure 1.4. Upper panel: annual mean surface salinity; middle panel: monthly mean surface salinity of January; lower panel: monthly mean surface salinity of July. Dataset from WOA13 (Reagan et al., 2013).

Figure 1 . 5

 15 Figure 1.5 Straits, seas and Indonesian Throughflow (ITF). NEC: North Equatorial Current; MC: Mindanao Current; NECC: North Equatorial Counter Current; SEC: South Equatorial Current; LC: Leeuwin Current. The number (for example 2.6Sv) represent the ITF flux at entrance and exits[START_REF] Gordon | Makassar Strait throughflow, 2004 to 2006[END_REF] Sprintall et al., 2009).

Figure 1 . 6

 16 Figure 1.6 Long term mean average monthly rate of precipitation in January (upper panel) and in July (lower panel). Data from: CMAP Precipitation data provided by the NOAA/OAR/ESRL PSD, Boulder, Colorado, USA, from their Web site at https://www.esrl.noaa.gov/psd

Figure 1 . 7

 17 Figure 1.7 Long term mean monthly wind direction in January (upper panel) and in July (lower panel). Data from: 20th Century Reanalysis V2c data provided by the NOAA/OAR/ESRL PSD, Boulder, Colorado, USA, from their Web site at https://www.esrl.noaa.gov/psd/

Figure 2 . 1

 21 Figure 2.1 Core locations of MD98-2165 and MD98-2166.

Figure 2 . 2

 22 Figure 2.2 Age model of MD98-2165. (a) Comparison of MD98-2165 benthic δ 18 O (Levi, 2003; Waelbroeck et al., 2006) with the global standard benthic δ 18 O stacks, LS16: 0 -150ka (Lorraine Lisiecki & Stern, 2016) and LR04: 150 -270ka (Lisiecki & Raymo, 2005). In the bottom, blue vertical bars indicate the AMS 14 C dates inTable 2.1 and green triangles represent

Figure 2

 2 Figure 2.3. Age model of MD98-2166. (a) Comparison of G. ruber δ 18 O between MD98-2165

Foraminiferal

  shell δ 18 O cal composition is controlled by temperature and δ 18 O sw composition of ambient seawater in which the foraminifers lived in, with the potential superimposition of species-dependent vital effects (Figure 2.4; Ravelo & Hillaire-Marcel, 2007). Due to the lighter weight of 16 O compared to 18 O and the impact this has the physics of water vapor production, during evaporation the water vapor is enriched in H 2 16 O relative to H 2 18 O compared to the initial liquid water. Thus, the δ 18 O values in water vapor, cloud droplets and precipitations are lower than in the initial seawater from which the vapor derived. During the transportation of water vapor, a Rayleigh distillation process takes place: H 2 18 O condensing more easily than H 2 16 O it is preferentially removed from cloud water vapor as precipitation, leaving the remaining water vapor increasingly depleted in δ 18 O with increased latitude and altitude (Ravelo & Hillaire-Marcel, 2007). Complex mechanisms of evaporation-transport-precipitation, superimposed upon the initial composition of ocean δ 18 O, result in specific, positive linear relationships between surface salinity and δ 18 O in the different ocean basins (LeGrande & Schmidt, 2006).

Figure 2 . 4

 24 Figure 2.4 Factors that affect the δ 18 O in foraminiferal calcite shell (modified after Ravelo &[START_REF] Ravelo | Chapter Eighteen The Use of Oxygen and Carbon Isotopes of Foraminifera in Paleoceanography[END_REF][START_REF] Hollstein | Variations in surface and thermocline conditions of the Western Pacific Warm Pool-Insights from Mg/Ca and stable isotopes in foraminifera tests[END_REF]

Figure 2 . 5

 25 Figure 2.5 Schematic process of reconstructing past local seawater δ 18 O local /salinity.(modified afterElderfield et al., 2012)

Figure 2 . 6

 26 Figure 2.6 Factors that affect Mg/Ca paleothermometer.

Figure 2 .

 2 Figure 2.12 Specification of standard solutions

Figure 1 .

 1 Figure 1. Locations of sediment cores discussed in this study; MD98-2165 (red dot, this study), water depth 2100m; MD98-2166 (red dot, this study), water depth 3875m; SO139-74KL (yellow dot) (Wang et al., 2018), water depth 1690m.

  displaying a clear glacial/interglacial structure in both cores with low Mg/Ca values during the glacial period increasing from around 20 ka Cal BP to reach a maximum at the onset of Holocene (~11 ka Cal BP). Core MD98-2165, which has a more complete Holocene section, shows decreasing Mg/Ca values towards a minimum centered around ~ 5 ka Cal BP and then increasing upward (~ 4.5 mmol.mol -1 at the core top). Overall, the G. ruber Mg/Ca records from both cores are rather similar in terms of absolute Mg/Ca values and relative glacial/early

  -1 (similar to glacial values) at the core top, with the most rapid drop in Mg/Ca occurring around 5 ka Cal BP. While the structure and values of Mg/Ca records were rather similar in both cores for G. ruber, the Mg/Ca records from cores MD98-2165 and MD98-2166 cores are significantly different for P. obliquiloculata. The deeper core MD98-2166 shows lower values than the shallower core MD98-2165 (Figure 3a-b) and does not reveal a strong glacial/interglacial structure, maximum Holocene Mg/Ca values (around 2 mmol.mol -1 )

Figure 3 .

 3 Figure 3. P. obliquiloculata Mg/Ca results from core MD98-2165 (a and c) and core MD98-2166 (b and d). (a) and (b), Mg/Ca ratios plotted versus age; the red dotted line shows Mg/Ca ratios obtained without the reductive step, the blue dotted line corresponds to Mg/Ca ratios obtained with the reductive step. (c) and (d), relative Mg/Ca-differences (%) (= (Mg/Ca without reductive step -Mg/Ca with reductive step ) / Mg/Ca without reductive step * 100). Averaged Mg/Ca-differences ± 1sigma uncertainty are displayed in red in (c) and (d), with n, the total number of samples. The horizontal dashed line (2 mmol/mol) is a visual guide to make it easier the comparison with Figure 4 (G. menardii).

(

  1s, n=53) in the deeper core MD98-2166 (Figure 4c-d). Such particularly strong variance in the deeper core is largely controlled by a few odd samples, which display much higher reductively-cleaned Mg/Ca values than their non-reductively cleaned aliquots (up to -30% difference; figure 4 d). Similar to what we observed for P. obliquiloculata in the deeper core MD98-2166, the combination of high variance in the Mg/Ca difference, and the low amplitude evolution of Mg/Ca along the core result in a very poor linear relationship between the reductively cleaned and non-reductively cleaned Mg/Ca values in MD98-2166 (R 2 =0.24).

Figure 4 .

 4 Figure 4. G. menardii Mg/Ca results from core MD98-2165 (a and c) and core MD98-2166 (b and d). (a) and (b), Mg/Ca ratios plotted versus age; the red dotted line shows Mg/Ca ratios obtained without the reductive step, the blue dotted line corresponds to Mg/Ca ratios obtained

  Figure 5 compares all our data in a X-Y plot, with reductively cleaned Mg/Ca values (Y) plotted versus non-reductively cleaned values (X). Since there is no evidence of a coherent evolution of A-B Mg/Ca differences along the cores (panels (c) and (d) in Figures 2, 3 and 4), such a

Figure 5 .

 5 Figure 5. Comparison of Mg/Ca ratios obtained with cleaning methods including and excluding the reductive step. Solid line, y=x; dashed line, linear best fits forced through the origin for samples from MD98-2165 (red), from MD98-2166 (blue) and all the data from the two sediment cores (green).

cleaning

  Mg-loss between the shallow core MD98-2165 (8.5% loss) and the deep core MD98-2166 (3.7%) as an indication that in situ dissolution, which has presumably removed some part of Mg-rich chamber walls (and, thus, has increased the relative contribution of the Mg-poor, dissolution resistant keel), significantly reduces the capacity of G. menardii to lose further Mg during reductive cleaning.Similar to what we observe for G. menardii, G. ruber data reveal a potential "pre-conditioning" effect of in-situ dissolution on the cleaning-related loss of Mg. Our data show that cleaningrelated offsets in Mg/Ca ratios are smaller for G. ruber (4.1%) in the deeper core MD98-2166 than from the shallower core MD98-2165 (6.2%). This difference is seen as statistically robust when running a t-test. This suggests that, similarly to what we observed for G. menardii, the tests of G. ruber from the deeper core, which have experienced more intense dissolution on the seafloor than the tests from the shallow core, are less prone to lose further Mg during reductive cleaning.

Chapter 4 :

 4 Northern Hemisphere forcing on Indo-Pacific Warm Pool hydrology change over the past 270, 000 years This chapter focuses on the hydrologic variations in the IPWP region and the low to high latitude teleconnections of climate changes through atmospheric circulation. In this chapter, we present a 270, 000-year record of local seawater δ 18 O composition-a sea surface salinity and hydrology indicator-estimated by combining Mg/Ca-SST and δ 18 O of G. ruber in core MD98-2165. This chapter addresses more particularly the relationship of hydrologic changes in the IPWP with respect to past variations of interhemispheric thermal gradient and AMOC dynamics.

Figure 1 .

 1 Figure 1. Locations of MD98-2165 and other ocean sediment cores discussed in the present paper, together with the locations of land cave records. ITF is the Indonesian Throughflow.The average, seasonal positions of ITCZ in January (austral summer) and July (austral winter) are shown by the faded red and blue lines. SO189-39KL(Mohtadi et al., 2014); SO139-74KL(Zhang et al., 2018); MD98-2165 (this study); MD98-2162(Visser et al., 2003). GeoB10069-3

Figure 2 .

 2 Figure 2. Sea surface salinity distribution at the surface (a) and at the depth of 100m (b). ITF is the Indonesian Throughflow. White dots and the red dot in (a) and (b) correspond to the main marine cores discussed in this study. 1:SO189-39KL(Mohtadi et al., 2014). 2:SO139-74KL (X.

3. 1

 1 Core MD98-2165 and δ 18 O stratigraphy Sediment core MD98-2165 was retrieved from a water depth of 2100 m during the Images IV cruise of the French R/F Marion Dufresne. Sediment core MD98-2165 is 42.3 m long; the upper 34.5 m are studied in the present paper. The core sediment consists of clay-carbonate mud with biogenetic and detrital silica. Pyrite is present all over the core. The sediment composition doesn't vary significantly from core top down to the bottom. Accelerated mass spectrometry

  . Past seawater δ 18 O sw can be calculated when one has an independent temperature proxy (e.g., Mg/Ca). This δ 18 O sw is a combination of global mean ocean δ 18 O (whose variations are controlled by changes in ice sheet volume) and local seawater δ 18 O. Ice volume corrected seawater δ 18 O sw-ivc changes reflect local or regional variations of the precipitation/evaporation budget and freshwater inputs (by river runoffs and lateral advection mechanisms linked to ocean circulation (e.g. LeGrande & Schmidt, 2011)). Thus, by estimating δ 18 O sw-ivc changes recorded in marine sediments, one can have access to past changes in seawater salinity associated to local/regional hydrological variations. Here, we use the toolkit PSU-Solver established by Thirumalai et al. (2016) to calculate past δ 18 O sw-ivc by combining G. ruber δ 18 O and Mg/Ca-T°. For the settings of PSU-solver, we chose the empirical calibration of Anand et al. (2003) for converting Mg/Ca data into temperatures (Mg/Ca = 0.38 exp 0.09 T) and we used the linear equation from Bemis et al. (1998) to derive δ 18 O sw (δ 18 O sw (VSMOW) = 0.27 + (SST(°C) -16.5)/4.8 + δ 18 O calcite (VPDB))

δ

  18 O change. The total SST increase during Termination III is ~3°C. In MIS7, SST were comparable between MIS 7e and 7c and slightly warmer than in the Holocene.During Termination II, SST shows a two-step increase. The first SST increase occurred at ~150 ka, leading changes in both planktonic and benthic δ 18 O records. The second step increase happened at ~ 136 ka, in phase with changes observed in the planktonic and benthic δ 18 O records. The SST reached its maximum at ~128ka, leading the minimum in the planktonic δ 18 O record. The total increase of SST from ~150 ka to ~128 ka is ~4.7 °C. Our record indicates that MIS 5e was the warmest period over the past 270 ka, with an average SST of ~29°C between 122 and 128 ka; this is about ~ 1 °C warmer than the early Holocene, and ~ 2°C warmer than during the MIS 5c.During MIS 2-4, SST was generally lower than 26 °C. During the Termination 1, the temperature increased by about 3.5°C, from a temperature of ~ 24.3°C during the LGM (~ 20 ka), to ~ 27.8°C during the early Holocene (~ 11 ka).4.2 δ 18 O sw and δ 18 O sw-ivcAs can be seen from Figure3c, reconstructed δ 18 O sw range between -0.42‰ and 1.69 ‰, showing the lowest value over the past 270 ka (~-0.2 ‰) during the middle to late Holocene (1-7 ka). δ 18 O sw from the top interval (~1.2ka) is ~ 0‰. In comparison, modern δ 18 O sw on the site of core MD98-2165 is ~0.1‰(LeGrande & Schmidt, 2006). On orbital time scale, δ 18 O sw of MD98-2165 varied in concert with global ocean mean δ 18 O sw (Figure3c). However, several discrepancies are observed, particularly during the Terminations, and as well as at ~60 ka and during the MIS 7c, clearly indicating the impact of local hydrological effects.

1

 1 , the δ 18 O sw-ivc reconstructed from core MD98-2165 vary between -0.4‰ and 0.94‰, and do not show an obvious glacial-interglacial pattern, glacial and interglacial values being rather similar (eg. δ 18 O sw-ivc is comparable during the warm MIS 5 and the colder MIS 4-2). The δ 18 O sw-ivc record shows a step-like evolution over the last 270 ka. Average values are heavy (around +0.3‰) in the interval 270-130 ka, then δ 18 O sw-ivc remain around a mean value of ~ +0.1‰ until termination 1. The δ 18 O sw-ivc is lower during the Holocene than during most of the other periods, with a mean value of -0.2‰ and peak values as low as -0.3‰. Values similar to the Holocene can be found, however, centered on ~ 70 ka. Within the older interval (270-130 ka), there is a clearly increasing trend during MIS6 from 190 to 130 ka. Several anomalous, positive excursions are superimposed on the long-term evolution. They occurred during the transition periods from cold glacials to warm interglacials: such as during termination IIIb (T-IIIb) at ~ 242 ka, during T-II at ~ 130 ka, during MIS4/MIS3 transition (~ 60 ka) and during the T-I, at ~ 16 and ~12 ka. In contrast, more negative excursions happened during the transitions from warm to cold stages: during MIS7e/7d at ~ 230 ka, during MIS7/MIS6 transition at ~ 190ka and during MIS5/4, at ~70ka. These observations suggest that local hydrology changes were most intense during periods of high amplitude climate change, towards a warmer or a cooler world. line indicate the AMS 14 C calendar dates and blue triangle represent tie points for correlating MD98-2165 benthic δ 18 O to LR04 and LS16 stack. Interpretation of the δ 18 O sw-ivc record: the salinity-δ 18 O relationship In situ precipitation and evaporation change, and as well as oceanic advection can affect the δ 18 O sw value of surface water. The strengthened (increased) evaporation (precipitation) increases (decreases) both salinity and δ 18 O sw . As a consequence, modern precipitation/evaporation processes coupled with the fractionation of water vapor during atmospheric transport result in a strong, positive link between surface ocean water δ 18 O sw and

  record with surface water δ 18 O sw-ivc records from other marine sites located in and out of the ITF pathways in the Indian Ocean side, and to land cave, stalagmite δ 18 O records, which are key archives of past, local rainfall changes. There are striking similarities between δ 18 O sw-ivc record of MD98-2165 and the Borneo cave stalagmite δ 18 O record. As can be seen in Figure 4a, all the anomaly enrichment of MD98-2165 δ 18 O sw-ivc , such as at 240 ka (T-IIIb), ~ 220 (T-IIIa), ~ 130 ka (T-II), ~ 60 ka (MIS4/3) and ~ 16ka (T-I), can find their counterparts in the Borneo cave δ 18 O records. The Borneo cave is located to the north of equator, near the upstream region of ITF (Figure 1). This cave record is interpreted as having recorded in situ (local) changes of precipitation over Borneo (Stacy A. Carolin et al., 2016); with low (high) rainfall resulting in the enrichment (depletion) in the Borneo cave δ 18 O. The strong positive correlation between the δ 18 O sw-ivc record of core MD98-2165 and the Borneo cave δ 18 O records readily suggests that past increase (decrease) in SSS at the site of Core MD98-2165 were associated to decrease (increase) of rainfall over Borneo.

Figure 4 .

 4 Figure 4. (a) Comparison of MD98-2165 δ 18 O sw-ivc and Borneo cave stalagmite δ 18 O (Carolin

Figure 5 .

 5 Figure 5. Comparison between δ 18 O sw-ivc record of MD98-2165 and other paleo-climate records over the last 30 ka. (a) GISP 2 ice core δ 18 O from Greenland (Grootes & Stuiver, 1997); YD, Younger Dryas; B/A, Bølling-Allerød; HS1, Heinrich stadial 1; LGM, last glacial maximum. (b) MD98-2165 Mg/Ca derived SST. (c) Relatively sea level(Waelbroeck et al.,

Figure 6 .

 6 Figure 6. δ 18 O sw-ivc enrichment event during Termination II. (a) MD98-2165 Mg/Ca derived SST. (b) Relatively sea level (Waelbroeck et al., 2002). (c) Borneo cave δ 18 O record (Carolin et al., 2016). (d) Comparison of δ 18 O sw-ivc records from different sediment cores in the eastern Indian Ocean (Figure 1 and 2).

  cave δ 18 O and the change of AMOC (Figure 7), we propose that these two curvesδ 18 O sw-ivc of MD98-2165 and Borneo cave δ 18 O -could be considered as an indicator of the change of interhemispheric thermal gradient. Those paleoclimate records, which are sensitive to the reorganization of atmosphere circulation driven by the change of interhemispheric thermal gradient should contain variations that have their counterparts in the MD98-2165 δ 18 O sw-ivc and Borneo cave δ 18 O records. One of such examples are the China cave stalagmite δ 18 O records. China caves located near the north boundary of modern ITCZ position (Figure 1), China cave stalagmite δ 18 O recorded the change of Asian Monsoon. Past meridional changes of the ITCZ have affected the strength of Asian monsoon. But, there is no consistent relationship between the original China cave δ 18 O and neither MD98-2165 δ 18 O sw-ivc record or the Borneo cave δ 18 O

  Cheng et al. (2016) have subtracted the northern hemisphere summer insolation signal (using July 21 insolation at 65°N as the main drive) from the China cave δ 18 O records (Figure7b and 7c). The residuals, called ∆δ 18 O, is interpreted as representing millennial-scale events of Asian monsoon as a response to the cold anomaly over the North Atlantic Ocean generated by the slowdown of AMOC(Cheng et al., 2016). The residual China cave signal, called ∆δ 18 O, in which orbitally-forced changes are smoothed out and 'millennial-scale' changes are enhanced, fits remarkably with both the Borneo cave δ 18 O record and our δ 18 O sw-ivc record from core MD98-2165 (Figure7c and 7d). As shown in Figure7cand 7d, all these three records exhibit strong similarities in both long-term and millennial scale change; positive excursion in the China cave ∆δ 18 O (interpreted as weak anomalies of suborbital Asian monsoon changes related to slowdown of AMOC) corresponding to increased Borneo cave δ 18 O (decreased rainfall) and increased δ 18 O sw-ivc of MD98-2165 (increased salinity). Thus, past changes in the average position of ITCZ correlated to AMOC changes (and thus, changes in the Conveyor Belt distribution of heat) appear to have modulated the strength of Asian summer monsoon (China cave ∆δ 18 O) and rainfall over the equatorial Indonesian seas (Borneo cave δ 18 O), which in turn modulated the input of salt in the Indian Ocean via surface water of the ITF (the δ 18 O sw-ivc of MD98-2165). Therefore, the δ 18 O sw-ivc of MD98-2165, Borneo cave δ 18 O and China cave ∆δ 18 O all reflect the change of the interhemispheric thermal gradient.

Figure 7 .

 7 Figure 7. (a) Comparison between MD98-2165 δ 18 O sw-ivc and North Atlantic 231 Pa/ 230 Th b (Böhm et al., 2015), AMOC is Atlantic Meridional Overturning Ciruculation. (b) Comparison between 12 July Insolation (65 o N)(Laskar et al., 2004) and China cave stalagmite δ 18 O(Cheng et al., 2016), AM is Asian Monsoon. (c) Comparison between China cave stalagmite ∆δ 18 O

MD98- 2165

 2165 δ 18 O sw-ivc varied coincidently with Borneo cave stalagmite δ 18 O record, and as well as with a millennial-scale China cave stalagmite ∆δ 18 O record. This suggest coherent hydroclimate changes over the Australian-Asian monsoon region. Large positive excursions of the MD98-2165 δ 18 O sw-ivc record occurred at each of the last three glacial terminations and at the transition between MIS 4/3, which is coeval with slowdown periods of AMOC. This indicate that eastern tropical Indian Ocean hydroclimate is impacted by Northern Hemisphere forcing through atmospheric teleconnections. In-phase oscillations recorded in the δ 18 O sw-ivc record of MD98-2165, the Borneo cave stalagmite δ 18 O record and the China cave stalagmite ∆δ18O record indicate that interhemispheric thermal gradient variations have always played a major role in tropical hydroclimate changes at millennial to orbital timescales. direction at surface (1981-2010). Data from: 20th Century Reanalysis V2c data provided by the NOAA/OAR/ESRL PSD, Boulder, Colorado, USA, from their Web site at https://www.esrl.noaa.gov/psd/. (e) and (f): Monthly mean sea surface temperature, data from WOA13(Locarnini et al., 2013). (g) and (h): Monthly mean sea surface salinity, data from WOA 13(Zweng et al., 2013). Red star indicates the location of sediment core MD98-2165.

Figure

  Figure S2. (a) Monthly mean sea surface temperature (SST), sea surface salinity (SSS) and precipitation rate at core site of MD98-2165 from the year of 2000 to 2015. SST and SSS data from Simple Ocean Data Assimilation (SODA) (Carton & Giese, 2008); precipitation data from CMAP Precipitation data provided by the NOAA/OAR/ESRL PSD, Boulder, Colorado, USA, from their Web site at https://www.esrl.noaa.gov/psd/. (b) Long-term monthly mean (2000-2015) SST, precipitation and SSS.

Figure S3 .

 S3 Figure S3. Testing the sensitivity of δ 18 O sw-ivc to the different selective of sea level curve. (a)Comparison of different sea level curve(Grant et al., 2012; Spratt & Lisiecki, 2016; Waelbroeck et al., 2002) and simulated sea level curve by MD98-2165 benthic-δ 18 O. (b)Comparison of δ 18 O sw-ivc obtained by using varied sea level cure in (a), pink shading region indicate 1 sigma error for the δ 18 O sw-ivc obtained by using the sea level curve of(Waelbroeck et

Figure 1 .

 1 Figure 1. Main pathways of Indonesian Throughflow (ITF) and associated surface currents (LC -Leeuwin Current; SEC -South Equatorial Current). Also indicated are the locations of sediment cores used in this study.

Figure 2 .

 2 Figure 2. Sea surface temperature distribution averaged over the January to March season (Austral summer, NW monsoon) and over the July to September season (Austral winter, SE monsoon). The maps show two different water depths: surface (0m) and subsurface (75m).

Figure 5 .

 5 Figure 5. (a) Monthly mean sea surface temperature (SST) at site MD98-2165 (red line) and SO139-74KL (blue line) from the year of 1985 to 2015 (Carton & Giese, 2008). (b) Monthly mean water temperature at depth 75m at site MD98-2165 (red line) and SO139-74KL (blue line) from the year 1985 to 2015 (Carton & Giese, 2008). (c) Nino 3.4 index; positive values indicate El Niño events; negative values indicate La Niña events (data source: http://www.cpc.ncep.noaa.gov/data/indices/oni.ascii.txt).

Figure 6 .

 6 Figure 6. Paleoclimate records from core MD98-2165. (a) G. ruber Mg/Ca derived SST (0-25ka, Levi et al., 2007). (b) P. obliquiloculata Mg/Ca derived TWT. (c) ΔT=SST-TWT. (d) Benthic foraminiferal δ 18 O of MD98-2165 (green line)(Levi, 2003; Levi et al., 2007;

  indicating a deeper thermocline depth during the colder MIS 8 than the Holocene and modern conditions.During Termination III, TWT increased at ~ 260 ka, leading the SST and benthic δ 18 O changes by ~ 10 thousand years. During MIS 7, TWT was higher during MIS 7e than the following warm sub-stages MIS 7c and 7a, whereas SST was comparable between MIS 7e, MIS 7c and MIS 7a. During MIS 6, both SST and TWT show relatively lower values.During Termination II, TWT showed an abrupt increase from ~136 ka and then reached nearly its interglacial values at ~133 ka, whereas SST other this interval only rose about 1/3 of the full deglaciation increase. After reaching the maximum at MIS 5e, TWT immediately decreased at ~ 128 ka, whereas SST sustained a high-value resulting in a clear plateau during MIS 5e until 120 ka, when SST showed the first decrease. Both SST and TWT display a long-term decreasing trend from MIS 5e towards to MIS 2. However, the difference, Mg/Ca-ΔT, does not exhibit any clear long-term trend in the interval ~120 to 30ka.From ~ 30 ka, Mg/Ca-ΔT displayed a decreasing trend till ~10ka, as a result of early and faster warming of the TWT relative to SST. During the Holocene, TWT decreased by about ~ 3°C, whereas SST changed less than 1°C. Consequently, Mg/Ca-ΔT reveals an increasing trend from 11 ka towards the present, with a marked increased trend over the last ~ 5ka.The above observations suggest that despite some similarities at the glacial-interglacial timescale, TWT changes do not mimic δ 18 O changes. Importantly, variations in TWT are not strongly linked to changes in glacial-interglacial boundary conditions during glacial terminations or inceptions as shown by the clear out-of-phase relationships between TWT and major changes in ice volume contained in the benthic δ 18 O record. The time-lag between these two proxies can reach up to ~ 10 ka (i.e. Termination III). Although the absolute value of such a time lag depends, obviously, on the quality of the age model (which rests upon a few tie-points only in the older part of the record), the fact that TWT and δ 18 O are out of phase during major glacial inceptions or terminations cannot be put into question since both records were obtained from the same sediment core.

Figure 7 .

 7 Figure 7. Spectral analysis of SST and TWT of core MD98-2165. (a) Original SST and TWT records; (b) spectral analysis result of original SST; (c) spectral analysis result of original TWT; (d) comparison of high frequency components of SST and TWT obtained by removing the >50-ka periodic components from the original SST and TWT (notched filtering (Gaussian)

146MD98- 2165 .

 2165 The absolute values of the TWT-SST offset vary with time (Figure8), although we do not know whether these apparent changes reflect potential problems in the age-model, or if they do reflect actual changes in lead/lag amplitude, which may have reflected different coupling between SST and TWT under different mean states of the climate and different forcings.

Figure 8 .

 8 Figure 8. Gaussian band-pass filter of SST (red) and TWT (blue) at orbital time scales. Band-

Figure 9 . 1

 91 Figure 9. Paleoclimate records comparison between core MD98-2165 and SO139-74KL

Figure 10 .

 10 Figure 10. (a) P. obliquiloculata Mg/Ca derived TWT of MD98-2165 and SO139-74KL (Wang et al., 2018). (b) ΔTWT=TWT(MD98-2165)-TWT(SO139-74KL), and JJA-Insolation:

Figure 11 .

 11 Figure 11. Comparison of MD98-2165 P. obliquiloculata Mg/Ca derived TWT and JJA-Insolation

Figure 12 .

 12 Figure 12. (a) Comparison P. obliquiloculata derived TWT between core MD98-2165, SO139-74KL(Wang et al., 2018) and MD01-2378(Xu et al., 2008), and JJA-Insolation gradient (20N-20S). (b) Relative contribution (%) of G. bulloides in the foraminiferal assemblages of coreGeoB10053-7, a proxy of upwelling activity and thus SE monsoon winds (large numbersindicate enhanced SE monsoon and vise versa(Mohtadi, Oppo, Steinke, et al., 2011)).

  intensity at ~10ka compared to today is also imprinted in the P. obliquiloculata Mg/Ca-derived TWT record of core SO139-74KL, whose lowest values are coeval with the G. bulloides peak in core GeoB10053-7. In this context of increased upwelling activity at ~ 10 ka compared to today, the P. obliquiloculata Mg/Ca-derived TWT record at site MD98-2165 shows maximum values at ~10ka, thus ~ 6°C higher than TWT site SO139-74KL. This suggests, contrary to what we observe today, the thermocline temperature signal at site MD98-2165 during the early Holocene was dominated by the stronger flux and/or warmer temperature of the ITF water. The TWT of MD98-2165 during the early Holocene was nearly ~3.5°C higher than temperatures recorder over the last ~1.5ka (Figure12). The higher P. obliquiloculata Mg/Ca derived TWT during the early Holocene could be also partly attributed to depressed deep nutrient upwelling water. Because decreased supply of nutrient would make the P. obliquiloculata flux rate tend to average on different season, resulting the P. obliquiloculata Mg/Ca to represent annual mean conditions during the Holocene, rather than the austral winter season condition as it does today.To have a depressed vertical upwelling during the increased offshore surface Ekman transport during early Holocene, one needs an increased water volume from the horizontal to compensate the losing of surface water. Assuming that temperature of thermocline ITF water was the same exiting Ombai Strait and the Timor Sea (away from the upwelling area), the difference of TWT between core MD98-2165 and core MD01-2378 should also be an indication of the relative strength of ITF flux relative to the upwelling intensity. We compared P. obliquiloculata Mg/Caderived TWT records between cores MD98-2165 and MD01-2378 in Figure12. Core top sediment studies indicate a 100 -120 m habitat depth for P. obliquiloculata in the Timor sea,

  slightly smaller peak glacials (i.e. lowest sea level of MIS 8, MIS 6 and MIS 2-4) compared to peak interglacials (i.e. highest sea levels of MIS 7 and MIS 5). This suggests slightly weaker SE monsoon winds during the lowest sea level periods. The relatively weak SE monsoon winds would increase P. obliquiloculata Mg/Ca derived TWT at site SO139-74KL by weakening the local upwelling, and would decrease the P. obliquiloculata Mg/Ca derived TWT at site MD98-2165 by weakening the ITF thermocline transport, thus generating small ΔTWT. In addition to the local SE winds related ITF volume change, the ITF temperature change as well as influence the TWT at site MD98-2165. We compare P. obliquilocualta Mg/Ca derived TWT between core MD98-2165 and MD06-3067 (Figure11)(Bolliet et al., 2011). MD06-3067 is located in western Pacific, in the inflow region of ITF, and thus P. obliquiloculata Mg/Ca TWT recorded the ITF source water temperature change. We found that TWT of MD06-3067 do not display a decreasing trend from 110 to 20ka which is observed in TWT record of MD98-2165. This result suggests the relatively lower TWT at site MD98-2165 during MIS 2-4 is caused by the increased local upwelling water effect as a result of reduced ITF transport. We infer that the relatively lower sea level and weak SE monsoon is responsible to the weak ITF transport during MIS2-4, MIS 6 and MIS8.

Figure 13 .

 13 Figure 13. Comparison of P.obliquiloculata derived TWT between MD98-2165, SO139-74KL

  sites (DTWT) appear to be nicely reflecting past changes in the relative upwelling and ITF strength. Spectral analysis reveals strong obliquity and precession-related components in the TWT record from core MD98-2165 and those orbital periods are also present in the DTWT record. Based on our understanding of the last 25 ka, our data suggest that a tight control of ITF relative strength by insolation always existed over the last 270 ka.On glacial-interglacial time scale, sea level changes may exert an effect on the TWT changes at site MD98-2165 by narrowing the ITF pathway or/and altering wind patterns (through modification of the land/sea distribution and albedo in the Indonesian Archipelago area). The change in ΔTWT between SO139-74KL and MD98-2165 is a result of change in SE monsoon winds and as well as the ITF heat transport, therefore can serve as an indicator to the ITF change in terms of both the heat and transporting volume. Take the most core top (~1.5ka) ΔTWT value (~1°C) as a standard, all the periods which ΔTWT is more than 1°C suggest a strong ITF characterized by increased transporting volume or/and increased heat. Based on this, we conclude that generally during MIS 8 and MIS 6 and MIS 4-2, ITF is weak, whereas during MIS 7, MIS 5 and MIS 1, ITF is relatively strong. This loss is equivalent to a 0.7 ± 0.9°C decrease in reconstructed temperatures. This loss results from the partial dissolution induced by the reductive cleaning step. But our data also indicate that the cleaning-induced Mg/Ca differences are species-dependent and reflect the preservation state of the tests. This explains why some inconsistencies were observed in the literature regarding the effects of cleaning. Although we evidenced obvious cleaning-induced Mg/Ca differences, there is no sign of significant changes in the relative Mg/Ca variations along a given core and for a given species when comparing the curves obtained by the two cleaning methods. In other words, relative seawater temperature changes reconstructed along sediment records (e.g. the temperature change between LGM and Holocene) are not affected by the choice of the cleaning method. Yet, as absolute seawater temperatures are affected, the paleoceagraphic community must develop clear strategies to construct coherent Mg/Ca-based temperatures. We suggest that the non-reductive cleaning approach should be preferentially used for building inter-laboratory homogeneous datasets of Mg/Ca-derived temperatures. The reductive step should be only included when severe oxide coating is observed (and, then, resulting Mg/Ca temperatures should be corrected for the cleaning offset using either a global correction factor (i.e. +6%) or, better, applying species specific correction factors. Although we have clearly established that in situ dissolution impact the cleaning-related Mg/Ca offset, we are aware that dissolution is a complex mechanism and that our two cores only gave us a partial view about the dissolution impact. Cores located at different water depths in the same areas are mandatory to fully explore the impact of weak to intense dissolution on Mg/Ca and Mg/Ca cleaning biases. At the moment, therefore, we assume that our best contribution regards the cleaning on well-preserved foraminifer shells of the three species studied.

6. 2 . 1

 21 Future work on core MD98-2165A -disentangling benthic temperature and δ 18 O sw signals in coreMD98-2165 In core MD98-2165 (retrieved at a water depth of 2100 m in the Indonesian Archipelago, eastern Indian Ocean), the G. ruber Mg/Ca-SST show good correlation with benthic foraminiferal δ 18 O without significant phase difference, especially during the main Terminations when Mg/Ca-SST varied synchronously with benthic foraminiferal δ 18 O. (Figure6.1) Benthic foraminiferal δ 18 O is a function of both bottom water temperature changes and bottom water δ 18 O sw , which chiefly represents global sea level changes. The modern bottom water at the site of core MD98-2165 is a mixture of North Atlantic deep water, circumpolar deep water and Red Sea deep water[START_REF] Fieux | Variability of the throughflow at its exit in the Indian Ocean: VARIABILITY OF THROUGHFLOW[END_REF] Waelbroeck et al., 2006). By reconstructing the bottom water temperature change, one can separate the ice-volume change component and temperature change component form the benthic foraminiferal δ 18 O. Reconstructing bottom water temperature from this high-sedimentation rate core, which contains particularly well-expressed benthic signals, would make it possible to look at the phase relationship between low-latitude surface temperature, the bottom water temperature and the global sea level change.

Figure 6 . 1 Figure 6 . 2

 6162 Figure 6.1 Comparison of G. ruber Mg/Ca-SST and benthic foraminiferal δ 18 O.

Figure 6 . 3

 63 Figure 6.3 Comprasion δ 18 O sw-ivc record between MD98-2165 (this thesis) and MD02-2594 (Martínez-Méndez et al., 2010).

Figure 6 . 4

 64 Figure 6.4 Comparison between MD98-2165 Mg/Ca-SST (this thesis) and Dome Fuji ∆T source[START_REF] Uemura | Asynchrony between Antarctic temperature and CO2 associated with obliquity over the past 720,000 years[END_REF].

Figure A2 .

 A2 Figure A2. Comparison of Elements/Ca ratios of G. ruber obtained with cleaning methods including and excluding the reductive step in core MD98-2165

Figure A15 .Figure A19 .

 A15A19 Figure A15. Comparison B/Ca ratios from different species cleaned without (CMA) and with (CMB) reductive step in sediment core MD98-2165

Table 2

 2

	.1 AMS 14 C dates and corresponding calendar ages of core MD98-2165*

Sampling depth (cm) Composite depth (cm) 14 C age ± !" (years BP) Calendar age (years BP) 68% (!") confidence interval (years BP)

	0	0	5250±100	--	--
	16	16	1650±40	--	--
	55	55	1150 ± 40	698	658 -728

Table 2 .2 δ

 2 18 O Tie points used for constructing MD98-2165 age model

Composite depth (cm) Age (ka) Type

	1259.53	61.21	δ 18 O Tie point
	1352.06	73.84	δ 18 O Tie point
	2009.68	132.50	δ 18 O Tie point
	2477.48	190.64	δ 18 O Tie point
	3124.89	246.07	δ 18 O Tie point

Table 2

 2

	.1 and green triangles represent

.1.1 Age model of core MD98-2166

  Globigerinoides ruber δ 18 O from core MD98-2166 to that of MD98-2165 with the new age model constructed in this thesis by using the LINAGE function from the ANALYSERIES software

	Sediment core MD98-2166 lacks AMS 14 C dates and a benthic δ 18 O record. Owing to the close
	location to core MD98-2165 (Figure 2.1), surface water physic-chemical properties (i.e.
	temperature and salinity, and thus δ 18 O of seawater) are similar today (Figure 2.3) and one can
	assume that past changes in hydrological characteristics were coherent and occurred
	synchronously at these two sites. Based on this assumption, age model for core MD98-2166 is
	established by tuning the planktonic,

Table 2

 2

	16.29	1.76	δ 18 O Tie point
	88.66	4.52	δ 18 O Tie point
	100.00	5.96	δ 18 O Tie point
	133.08	9.23	δ 18 O Tie point
	412.74	20.60	δ 18 O Tie point
	778.89	54.72	δ 18 O Tie point
	1058.80	67.68	δ 18 O Tie point
	1160.49	79.75	δ 18 O Tie point
	1410.08	109.55	δ 18 O Tie point
	1489.97	126.71	δ 18 O Tie point
	1550.03	136.50	δ 18 O Tie point
	1872.19	164.48	δ 18 O Tie point
	2075.28	204.57	δ 18 O Tie point
	2162.97	220.73	δ 18 O Tie point
	2236.33	225.95	δ 18 O Tie point
	2330.86	239.32	δ 18 O Tie point

.3. δ 18 O Tie points used for correlating MD98-2166 to MD98-2165 Composite depth (cm) Age (ka) Type

Table 2 .5 Elements measurement setting

 2

	Elements	Dwell time (µsec)	Attenuation mode
	Li7	20000	None
	B11	20000	None
	Mg24	20000	Medium
	Ca43	20000	Medium
	Sr88	10000	High
	Cd112	20000	None
	Ba137	20000	None
	U238	20000	Medium
	Al27	10000	Medium
	Mn55	10000	Medium
	Fe57	10000	Medium

Table 2 .6 The relative standard deviation (RSD) of measuring Elements/Ca

 2

	Elements/Ca	RSD (n=144)
	Li7/Ca43	2.3%
	B11/Ca43	8.2%
	Mg24/Ca43	1.4%
	Sr88/Ca43	1.8%
	Cd112/Ca43	4.3%
	Ba137/Ca43	3.3%
	Al27/Ca43	5.5%
	Mn55/Ca43	5.6%
	Fe57/Ca43	7.0%
	U238/Ca43	4.7%

  Difficulties in drawing consistent conclusions from previous works come from the fact that no comprehensive study of cleaning effects has yet combined all the following characteristics:(1) the two cleaning protocols should be applied on the same samples, within the same laboratory, to limit other potential sources of bias, (2) the cleaning protocols should be applied to different planktonic species picked from the same samples in order to test for inter-species differences, (3) the potential impact of the preservation state of foraminiferal tests should be addressed by comparing sedimentary material retrieved at different water depths, (4) the study should include a downcore evaluation of Mg/Ca offsets related to cleaning methods. Such a downcore comparison is mandatory if one wants to decipher whether the cleaning offsets may vary through time or depth, along the same sedimentary record, as a result of changes in the abundance of oxides (reflecting past variations in bottom water

	cleaning protocol. To make things even more puzzling, Johnstone et al. (2016) recently reported
	no Mg/Ca change for P. obliquiloculata after reductive cleaning, and relative Mg/Ca losses of
	~4% for G. ruber, ~10% for Neogloboquadrina dutertrei and ~5% for Globigerinoides
	sacculifer.

observed an average reduction of ~15% in Mg/Ca ratio obtained from foraminiferal tests cleaned with a reductive step compared to those cleaned without it.

Barker et al. (2003)

had previously reported a similar 10-15% shift in Mg/Ca between the two cleaning approaches.

Xu et al. (2006)

reported, however, much smaller relative Mg/Ca differences and suggested that Mg-loss could be species dependent, with a ~6.6% Mg/Ca relative decrease for G. ruber, and a smaller ~ 1.8% loss for Pulleniatina obliquiloculata.

Fan et al. (2018)

, on the other hand, recently reported a stronger ~ 6% Mg/Ca decrease for P. obliquiloculata when a reductive step was added to the As can be seen for these few examples, published data are confusing, fuelling a never-ending debate about cleaning methodologies and how to compare or even combine Mg/Ca results obtained from different laboratories. oxygenation), past changes in calcite preservation conditions at the seafloor, and/or diagenetic effects. In addition, downcore studies provide a large number of samples, making it possible to derive more statistically robust evaluations of cleaning effects.

Ice volume corrected seawater δ 18 O sw-ivc

  ) along a one-year period is ±1.4% (n=144, 1s). Mn/Ca, Fe/Ca and Al/Ca were measured together with Mg/Ca for detecting clay contamination and post-depositional Mn-rich oxide coatings. No relationship was observed between Mg/Ca and Mn/Ca, Fe/Ca or Al/Ca variations. Thus, no Mg/Ca data was rejected from our dataset. SST were calculated from

	Composite depth (cm) 14 C age ± !" (years BP) 0 5250±100 mmol/molMg/Ca of G. ruber through the calibration equation established by (Anand et al., 2003): Mg/Ca Calendar age (years BP) 68% ( !" ) confidence interval (years BP) ----
	16 = 0.38 exp 0.09 T. The errors in SST reconstruction is estimated by propagating the error 1650±40 ----
		55	1150 ± 40	698	658 -728
	135 introduced by Mg/Ca measurements and Mg/Ca temperature calibration. The resulting error is 2210 ± 40 1810 1756 -1866
	213 260 on average about ±1.2°C (1s) 3210 ± 40 3910 ± 90	3016 3891	2943 -3077 3760 -4020
		320	4740 ± 110	5009	4843 -5134
		460	7500 ± 110	7962	7841 -8079
	3.4	560	10340 ± 120	11446	11208 -11626
		600	11430 ± 130	12910	12761 -13049
		620	12100 ± 130	13561	13414 -13699
		640	13040 ± 150	14930	14656 -15268
		656	13590 ± 140	15826	15625 -16058
		664	14050 ± 120	16464	16257 -16643
		674	14360 ± 140	16923	16679 -17146
		680	15080 ± 160	17856	17650 -18041
		690	15660 ± 130	18516	18375 -18666
		698	16030 ± 140	18881	18728 -19018
		740	19780 ± 170	23329	23093 -23548
		774	22330 ± 220	26170	25933 -26477
					. After cleaning, each
	sample was dissolved in 300 µl of 0.075M HNO 3 and then centrifuged for removing
	undissolved particulate impurities. The supernatant solution was transferred to a new acid-
	cleaned 500 µl centrifuge tube. Mg/Ca ratios were measured on an inductively coupled plasma
	mass spectrometer (ICP-MS, Analytik Jena PlasmaQuant MS Elite). Aliquots of samples were
	first analyzed for Ca concentrations and the remaining solutions were diluted to reach a Ca
	concentration of ~100 ppm in order to minimize possible matrix effects. A standard solution
	was analyzed every 3 samples for monitoring the instrument behavior and correcting the

samples for potential drifts and instrumental instabilities. The long-term reproducibility of Mg/Ca measurements obtained by replicating analyses of a standard solution (Mg/Ca=3.69

  ICP-MS, Analytik Jena PlasmaQuant MS Elite), and as well as for Al/Ca, Mn/Ca and Fe/Ca for examining the cleaning efficiency on clay contamination and postdepositional Mn-rich oxide coatings. Aliquots of samples were first analyzed for Ca concentrations and the remaining solutions were diluted to reach a Ca concentration around 100 ppm for all the samples in order to minimize possible matrix effects. A drift correction standard No relationship was observed between Mg/Ca and Mn/Ca, Fe/Ca or Al/Ca variations.

	Composite depth (cm) 14 C age ± !" (years BP) 55 1150 ± 40 135 2210 ± 40 213 3210 ± 40 plasma mass spectrometry (solution was analyzed every 3 samples for monitoring the instrument variability and correcting Calendar age (years BP) 68% ( !" ) confidence interval (years BP) 698 658 -728 1810 1756 -1866 3016 2943 -3077
	260 samples after the analysis. The long-term reproducibility of Mg/Ca measurements obtained by 3910 ± 90 3891 3760 -4020
	320	4740 ± 110	5009	4843 -5134
	460 replicating analyses of a standard solution (Mg/Ca=3.69 mmol/mol) is ±1.4% (1s). This 7500 ± 110 7962 7841 -8079
	560 600 estimate is based on n=144 analyses performed during 24 working sessions over a one-year 10340 ± 120 11446 11208 -11626 11430 ± 130 12910 12761 -13049
	620 640 period.	12100 ± 130 13040 ± 150	13561 14930	13414 -13699 14656 -15268
	656	13590 ± 140	15826	15625 -16058
	664	14050 ± 120	16464	16257 -16643
	674	14360 ± 140	16923	16679 -17146
	680	15080 ± 160	17856	17650 -18041
	690	15660 ± 130	18516	18375 -18666
	698	16030 ± 140	18881	18728 -19018
	740	19780 ± 170	23329	23093 -23548
	774	22330 ± 220	26170	25933 -26477

Table 2 .

 2 Table of sediment records discussed in this study showing the core locations, the foraminiferal species used, the size fraction in which they were picked and the cleaning methods used for Mg/Ca measurements.

	Core ID	Lat./Long. (water depth)	Species (size)	Time	Cleaning* Reference
				period
	MD98-2165	9.65 S, 118.34 E (2100m)	G. ruber (250 -315 µm)	0 -30 ka	NRC	Levi et al., 2007
			G. ruber (250 -350 µm)	30-270 ka	NRC	This study
			P. obliquiloculata (250 -350 µm)	0 -270 ka	NRC	This study
	SO139-74KL 6.54 S, 103.83 E (1690m)	G. ruber (250 -350 µm)	0 -300 ka	RC	Wang et al.,
						2018
			P. obliquiloculata (350-450 µm)	0 -300 ka	RC	Wang et al.,
						2018
	MD01-2378	13.08 S, 121.79 E (1783m)	P. obliquiloculata (250 -315 µm)	0-25 ka	RC	Xu et al., 2008
			G. ruber (250 -315 µm)	0 -160 ka	RC	Bolliet et al.,
						2011
	MD06-3067	6.51 N, 126.50 E (1575m)	P. obliquiloculata (250 -315 µm)	0 -160 ka	RC	Bolliet et al.,
						2011
	*NRC represents non-reductively cleaning methods, RC represents reductively cleaning methods.

Table A1 . The average Elements/Ca ratios differences between two cleaning methods

 A1 Elements/Ca Difference (%) = (Elements/Ca without reductive step -Elements/Ca with reductive step ) / Elements/Ca without reductive step * 100%; P.obliq. is short of P. obliquiloculata.

			Li/Ca	B/Ca	Mg/Ca Sr/Ca Cd/Ca Ba/Ca Mn/Ca Fe/Ca	U/Ca
	MD98-2165	G.ruber P.obliq. G.menardii	1.3% 1.3% 5.8%	-6.3% -2.2% 1.0%	6.2% 4.4% 8.5%	-1.2% -0.3% 4.4%	85.6% 65.7% 63.8% 75.2% 7.8% 60.3% 74.4% 16.3% 58.0%	4.7% 2.9% 5.0%	------
	MD98-21666	G.ruber P.obliq. G.menardii	1.0% 3.9% 2.4%	-3.3% -2.1% -1.7%	4.1% 6.3% 3.7%	0.7% 2.2% 0.8%	87.0% 46.0% 59.3% 83.1% 20.6% 66.0% 80.9% 20.2% 56.1%	1.0% 0.1% 0.8%	76.9% 72.5% 70.4%

*

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Acknowledgements

I am grateful for the last four years, during which I have been living in France and working on my PhD project. Many people accompanied me through out that time.

I thank my advisors, Franck Bassinot and Sophie Sepulcre, for their time, patience, instructions, supports and etc. I especially thank Franck, who offer me the PhD position. Paleoceanography study is full of uncertainties and noises, thank Franck for teaching me how to fish in the ocean noise.

I thank Isguder Gulay for her help on identifying and picking foraminifers. I thank Helene Rebaubier and Haowen Dang for their collaboration on working with ICP-MS. I thank Claire Waelbroeck for providing me the oxygen isotope data. I thank William Gray for his helpful discussion. I thank Elisabeth Michel for offering the opportunity working on board of R/V L'Atalante. I thank Herve Guillou for his many kindly help. I would like to thank all the

Figure 6. (a) Comparison of G. ruber Mg/Ca-SST with using different cleaning methods in core MD98-2165, and cleaning-offset correction; SST is the short of sea surface temperature; TWT is the short of thermocline water temperature; NRC represents non-reductively cleaned samples and RC represent reductively cleaned samples. (b) Comparison of P. obliquiloculata Mg/Ca-TWT in core MD98-2165. (c)(d)(e) Comparison of G. ruber Mg/Ca-SST between core MD98-2165 and SO139-74KL (Wang et al., 2018) cleaned with and without reductive step and cleaning-offset correction.

Figure 3. Paleoclimate records of core MD98-2165. (a) Planktonic δ 18 O of G.ruber (Levi, 2003). (b) G. ruber Mg/Ca derived SST (0-25ka, Levi et al., 2007). (c) δ 18 O sw of seawater of MD98-2165 (blue line) and sea-level equivalent δ 18 O (black line) (Waelbroeck et al., 2002). ( d) Ice volume change corrected δ 18 O sw-ivc record, pink shading region indicate one sigma error.

(e) Green curve is benthic δ 18 O of core MD98-2165 (Levi, 2003; Waelbroeck et al., 2006); black curve is global benthic δ 18 O stack, 0-150 ka is from LS16 (Lisiecki & Stern, 2016) and 150-270 ka from LR04 (Lisiecki & Raymo, 2005)

; T-I, T-II, T-III represent Terminations; MIS is marine isotope stage, sub-stage of MIS5(5a-5e) and sub-stage of MIS7 (7a-7e); red vertical

Figure 8. Spectral analysis results of MD98-2165 δ 18 O sw-ivc and China cave ∆δ 18 O records. (a) Comparison between original δ 18 O sw-ivc of MD98-2165 (black line) and residuals after removing the >50-ka cycle components. (b) Spectral analysis results of MD98-2165 δ 18 O sw-ivc (original) and δ 18 O sw-ivc (<50ka) in (a). (c) Comparison between MD98-2165 δ 18 O sw-ivc (<50ka) and China cave ∆δ 18 O. (d) Spectral analysis results of MD98-2165 δ 18 O sw-ivc (<50ka) and China cave ∆δ 18 O as displayed in (c). (e) and (f) Gaussian band-pass filter of China cave ∆δ 18 O (black line) and MD98-2165 δ 18 O sw-ivc (red line) at precession and obliquity band, Band-pass filtering was performed by using software of AnalySeries 2.0.8 (Paillard et al., 1996), the corresponding obliquity component were filtered out from China cave ∆δ 18 O (black line) and MD98-2165 δ 18 O sw-ivc at frequency 0.026±0.004 ka -1 ; The corresponding precession component were filtered out from China cave ∆δ 18 O (black line) and MD98-2165 δ 18 O sw-ivc at frequency 0.044±0.004 ka -1 .

Abstract

The aim of this study is to better assess the impact of cleaning methods on the Mg/Ca of dissolution-resistant and dissolution-sensitive planktonic foraminifers showing different preservation states. For this purpose, we performed a downcore comparison of Mg/Ca ratios obtained using reductive and non-reductive cleaning methods applied to Globigerinoides ruber, Pulleniatina obliquiloculata and Globorotalia menardii picked along two nearby sediment cores retrieved at 2100m and 3875m in the eastern equatorial Indian Ocean. Combining all data together, we observe an average loss in Mg/Ca of ~ 6 ± 7% (1s, n=328) when a reductive step is included in the cleaning method. This loss is equivalent to a 0.7 ± 0.9°C decrease in reconstructed temperatures. This loss results from the partial dissolution induced by the reductive cleaning step. Our data reveal a rather large scattering between reductively cleaned and non-reductively cleaned Mg/Ca for P. obliquiloculata and, more particularly, G. menardii aliquots. We interpret this scattering as chiefly resulting from inhomogeneous Mg distribution within the tests of these two species, making it difficult to split the fragments into perfectly homogeneous aliquots for comparing the two cleaning methods. Despite this scattering, our results show statistically robust discrepancies between the three species and the two core sites, suggesting that cleaning-induced Mg/Ca differences (i) are species-dependent and (ii) reflect the preservation state of the tests. We conclude that the non-reductive cleaning approach should be preferred for building inter-laboratory homogeneous datasets of Mg/Ca-derived temperatures, whereas the reductive step should be only included when severe oxide coating is observed. SSS at the exit of the Timor Sea. During these HS1 and YD events, the Borneo and Flores records are out-of-phase, with the Flores cave δ 18 O record showing depleted values, which indicate that local rainfall, in the vicinity of core MD98-2165 site was high at the time when SSS was high. At present, about 70% of mean annual rainfall (~1200mm) at the site of Flores cave occur during the austral summer wet season from December to March (Ayliffe et al., 2013).

Northern

The depleted Flores cave δ 18 O values during YD and HS1 suggest increased austral summer rainfall. As already mentioned above, the out-of-phase relationship between the Flores and Borneo records is a direct result of past, meridional shift of the main position of ITCZ (Ayliffe et al., 2013;Denniston et al., 2017;Krause et al., 2019). Our observations support, therefore, the hypothesis that the δ 18 O (and SSS) variations at the site of core MD98-2165 do not chiefly reflect local precipitation changes, but rather lateral advection and the salinity of surface ITF waters.

In the Indonesian region G. ruber shell chemistry (i.e., here is δ 18 O sw-ivc ) represent annual mean surface conditions (Mohtadi et al., 2009). The fact that surface δ 18 O sw-ivc record obtained by analyzing G. ruber shells can smooth out seasonal variations could be one of the reasons for the decoupling between land cave (Flores cave) and seawater (MD98-2165) δ 18 O change during YD and HS1. But it is not the main element. Seawater δ 18 O reconstructed from a sediment core near Borneo show striking covariance with the Borneo cave δ 18 O record (Zhang et al., 2018), as do the MD98-2165 record (this study) and a record from core SO189-39KL, off western Sumatra (Mohtadi et al., 2014; Figure 5). It is important to notice, however, that precipitation over Borneo (~ 5000mm) is much higher today than over Flores (1200mm), and does not show a marked seasonal change (Carolin et al., 2016). Our results suggest, therefore, that the strongest imprint over surface ITF δ 18 O (and SSS) in the Indonesian Archipelago comes from heavy precipitations taking place in the central part of Indonesian Seas around the equator, allowing to promote a clear coupling between the on land cave Borneo record, and ocean seawater δ 18 O records. Rainfall changes, which impacted the δ 18 O sw in Borneo and in equatorial Indonesian region would be transported laterally to the core site of MD98-2165 by lateral advection of the surface ITF (Figure 1). These results strengthen our initial hypothesis that past changes in MD98-2165 δ 18 O sw-ivc are chiefly controlled by variations in ITF strength and salinity rather than in situ changes in precipitation and evaporation. The sea surface temperatures (SST) and thermocline water temperature (TWT) were reconstructed by the foraminiferal Mg/Ca from the mixing layer species G. ruber and thermocline water species P. obliquiloculata respectively. Before using Mg/Ca paleothermometry, the impact of cleaning methods (with or without a reductive step to remove oxides) on planktonic foraminiferal Mg/Ca analyses were evaluated. The sea surface salinity (SSS) was evaluated by reconstructing δ 18 O of seawater through the combination of Mg/Ca-SST and δ 18 O of G. ruber.

Supplementary Figures

By analyzing these new data and combining other paleoclimate records in the IPWP, the main conclusions from this work are:

(1) In foraminiferal Mg/Ca analysis, including a reductive step in cleaning methods leads to an average loss in Mg/Ca of ~ 6 ± 7% (1s, n=328), when combining results from the three species G. ruber, P. obliquiloculata and G. menardii from both core MD98-2165 and MD98-2166).

All these elements can be explained by past changes in the dynamics of the ITCZ. Large positive excursions of the MD98-2165 δ 18 O sw-ivc record occurred at each of the last three glacial terminations and at the transition between MIS 4/3, which are coeval with slowdown periods of AMOC as revealed by a recent Pa/Th record. This suggests that rapid changes in the eastern tropical Indian Ocean hydroclimate are impacted by Northern Hemisphere forcing, which affects the dynamics of ITCZ. In-phase oscillations recorded in the δ 18 O sw-ivc record of MD98-2165, the Borneo cave stalagmite δ 18 O record and the China cave stalagmite ∆δ18O record indicate that interhemispheric thermal gradient variations have always played a major role in tropical hydroclimate changes at millennial to orbital timescales.

(3) We showed that changes in P. obliquiloculata Mg/Ca derived TWT at site MD98-2165 reflect the balance between the strength of ITF and the local southeast (SE) monsoon winds driven upwelling intensity. At millennial to orbital timescales, enhanced SE monsoon winds lead to strengthened ITF transport. On glacial-interglacial scale, the ITF is modulated by the sea level change, with showing weakened ITF linked to lower sea level. Our results shown that generally ITF was weak during cold stage MIS 8, MIS 6 and MIS 2-4, whereas ITF was stronger MIS 7, MIS 5 and during MIS 1. Especially, the ITF transport in both flux and heat is at maximum peak during Northern Hemisphere maximum summer insolation.

Perspectives and future work

During the course of my stay at LSCE, I've spent a large amount of work and time to set up the analytical protocol on the new ICP-MS (which arrived in April, 2016 after the beginning of my PhD), run tests on this new machine to understand its mode of variability and potential biases and set up the reductive cleaning procedure (in the clean laboratory of Geops, Orsay), which had never been used at LSCE for foraminiferal Mg/Ca analyses before. Due to the large amount of time and efforts devoted to this analytical aspect of the work (and the participation to an oceanographic cruise in the Atlantic cruise), I may have felt frustrated sometime not to have enough time to explore all the results I gathered and, in particular, the wide range of minor and trace elements that I analyzed together with Mg/Ca. I've also envisioned potential issues to go further into the interpretation of the Mg/Ca-derived temperature records that I obtained.

Following, are a few issues that I think worth developing in the future and that could be a focus of research that I would be most willing to pursue during post-doctoral research, back in China.