Xavier Carton
email: xcarton@univ-brest.fr

P L'hegaret

Rémy Baraille

Pierre L'hégaret
email: pierre.lhegaret@univ-brest.fr

Léo Lacour

Guillaume Roullet

Stéphanie Corréard

Jin-Song Von

Storch P L'hégaret

R Duarte

C Vic

D Ciani

Pierre L'h Egaret

Isabel Ambar

Claire M Enesguen

Bach Lien Hua

Laurent Ch Erubin

Ana Aguiar

Bernard Le Cann

Nathalie Daniault

Nuno Serra

Mesoscale variability of water masses in the Arabian Sea as revealed by ARGO floats

Keywords:

Je souhaite remercier Fabien Durand et Gilles Reverdin qui ont généreusement accepté d'être rapporteurs de ma thèse et dont les remarques m'ont permis d'augmenter la qualité de ce manuscrit.

Je tiens également à remercier Rosemary Morrow, Carole Nahum et Laurent Mémery pour leurs participations à mon jury.

Je souhaite exprimer mes profonds remerciements à Xavier Carton, dont le soutien, les conseils et les encouragements m'ont permis de présenter cette étude.

The Indian Ocean is the third largest oceanic area of the world, however, it is quite di↵erent from the Atlantic and Pacific Ocean. A first noteworthy variation is the presence of the continents bordering the ocean in the northern hemisphere. This presence, due to the intense geophysical activity in the region, sets one of the strongest land/air/sea interaction, the monsoons. Along the year, with two reversals, the winds blow steadily over the Indian Ocean, bringing heavy precipitations over the lands and forcing the circulation of the ocean.

In the northern hemisphere, the complex tectonic plates interactions has formed several basins, subject to the monsoon, each one with specific water mass characteristics and circulations. West of India, North of the Equator, the Indian Ocean is comprised by five hydrological basins : the Arabian Sea, extending in two gulfs, the Gulf of Aden and the Gulf/Sea of Oman, each opening on two semi-closed, intense evaporative basins producing highly saline waters, the Red Sea and the Arabian/Persian Gulf.

General Introduction

Since the XV th century, with Vasco de Gamma opening the first direct sea lane between Europe and India, the northwestern Indian Ocean has been a region of high interest, through trades. The Persian Gulf nations are presently the principal oil and naturals gas producers and possess the first reservoirs of crude oil, thus making it a region of high geopolitic interest.

Nevertheless, the oceanic expeditions have been scattered through this basin, until 1959-1965 when the United Nations launched the International Indian Ocean Expedition aiming at a systematic oceanographic investigation of the Indian Ocean. However, expeditions in these basins concerned only one specific region or one specific monsoon. Thus, some phenomena are well studied and documented, such as the water mass formation, their exchange through the straits or the global circulation of the Indian Ocean.

The monsoon dominates the dynamics of the Indian Ocean. In the Arabian Sea, down to the Equator, the winds, through the Ekman transport, chase waters towards the coasts thus creating a basin wide cyclonic circulation in winter. In summer these waters and chase o↵ coast, creating strong upwellings and an anticyclonic circulation in summer over the Arabian Sea. Along the coasts, principally of India, Oman, Yemen and Somalia, and along the equator, this circulation is expressed via currents, reversing depending of the seasons. Nevertheless, this circulation is complex, particularly near the basins opened through the tectonic plate dynamics, the Gulf Of Aden and the Sea of Oman and a study from a mesoscale perspective of the dynamics in this region is of interest.

Nowadays, ARGO floats1 are and have been seeded around the globe The floats have sensors recording the water characteristics at depth, thus providing a continuous sampling of the oceans. Over the last 12 years, about 300 ARGO floats were launched in the Arabian Sea and the adjacent basins, representing over 20 000 profiles. Note that these floats also provide implicit informations on the water circulation by cycling with a 5-10 days period.

Thus, the first aim of this thesis is to study, for each season, the circulation of the di↵erent water masses in the North-West Indian Ocean from a mesoscale perspective. The joint use of observational data from the ARGO floats, climatology, satellite data and reanalysis provides us the first approach of this problem in the Arabian Sea and two adjacent gulfs, the Gulf of Aden and the Sea of Oman. The surface circulation appears to be dominated by large mesoscale eddies, with a strong vertical impact on the water masses at depth, particularly on the highly saline water produced in the Persian Gulf and in the Red Sea.

From these observations, an emphasis is then put on the mesoscale eddies, from their formation and evolution to their impact on the Persian Gulf Water, over the di↵erent seasons. For that, analysis tools are used on both the observational data and outputs from a numerical model (the primitive equations model HYCOM). Some processes are also studied with idealized models. Thus, this provides us a description of the sub-basins, the 3-D structures of the mesoscale features and their seasonal variability.

Finally, a dataset from the Phys-Indien 2011 experiment, providing higher resolution in-situ data around the Arabian Peninsula, in March 2011. The measurements display the interactions between the eddies and the water masses in the Arabian Sea and Sea of Oman. An emphasis is put on the characteristics and trajectory of the Persian Gulf outflow in the Sea of Oman during this period from the mesoscale to the submesoscale, where this water mass is found under the form of lenses of filaments.

Thus, the main questions addressed in this thesis are : -Which features dominates the mesoscale circulation in the Arabian Sea ? -How do they form and what are their structures ? -How does the monsoon seasonality impact them ? -What is the interaction between these mesoscale structures and the water masses of the region ? -What is the evolution of the Persian Gulf outflow pathway in the Sea of Oman throughout a whole year ? -What are the characteristics of the highly saline waters in the Sea of Oman and Arabian Sea, from a mesoscale and submesoscale perspective ?

The thesis is organized as follows. In a first chapter, a wider presentation of the region is given, with a description of the geography, climate and historical oceanography. Then an emphasis is put on the water mass formation and circulation in each basin. The second chapter focuses on the di↵erent datasets used during the studies and on their processing. The Phys-Indien 2011 experiment, the devices used, the locations of the measurements, the biases and the processing are detailed. Furthermore, descriptions of the numerical studies used over the di↵erent parts of the work are given, with an emphasis on the HYCOM model. The third chapter qualitatively and quantitatively describes ARGO float measurements in the Arabian Sea in the Gulf of Aden, in particular to track Red Sea Water and Persian Gulf Water, highly saline waters that equilibrate in the northern Indian Ocean at 600 and 200 m depth respectively. Their trajectories and measurements are linked with the mesoscale circulation of these basins. This chapter is centred on an article published in Ocean Science on March 2012. The fourth chapter also used ARGO floats, but in the Sea of Oman. The onset of a dipole is observed recurrently during the springs. A dynamical analysis of this structure is given, and its interaction with the wind forcing and propagating Rossby waves is analysed. The core of this chapter is an article published in Ocean Dynamics in 2013. The fifth chapter aims at describing the global circulations in the Arabian Sea through each seasons via long term observations and numerical simulations with HYCOM. The mesoscale features are described over their life cycles, with an emphasis on their impact on the Persian Gulf outflow. This chapter is presented as an article, submitted to Ocean Science early January 2015. The sixth chapter details the measurements from the Phys-Indien 2011 campaign and compares the circulation for this year with the average circulation presented in the previous chapter. With the higher resolution of this dataset, an analysis of the characteristics of the Persian Gulf Water, and of the surrounding water masses is given at the mesoscale and submesoscale. Then perspectives are given in a conclusion. This manuscript includes articles written along the thesis, in a chronological order, as the studies progressed. Each one is directly followed by a section describing the complementary work related to these papers.

In an appendix, an article published in the Journal of Geophysical Research in July 2014 is added to this manuscript. This study was done during the period of this thesis, about the Gulf of Cadiz, with a joint use of three datasets from di↵erent experiments, General Introduction satellite measurements and numerical modelling to describe the merging of two mesoscale dipoles, containing Mediterranean Water. This paper was added here because it also concerns the mesoscale dynamics of a warm and salty outflow, though in a di↵erent region than for the rest of the thesis. It also used similar tools, and in that respect, shows relation with the main body of the thesis. Finally it took several months to complete and thus represents part of the doctoral work.

Topography

The northwestern Indian Ocean is a region of strong geophysical activity with the interaction of no less than five di↵erent plates : the African and Somalia plates, the Indian plate, the Arabian plate and the Eurasian plate. From a geographical point of view, this includes the Middle-East, around the Arabian Peninsula, the Horn of Africa and the western Indian subcontinent, composed of India and Pakistan. The oceanic crust of the Arabian plate contains the eastern Red Sea, the northern Gulf of Aden, the northwestern Arabian Sea, extending to the southwestern Sea of Oman and the Persian Gulf (see map on figure 1.1).

We present hereafter each plate interaction with the Arabian plate and their geographical formations. The African plate moves away from the Arabian plate, this spreading formed the Red Sea Rift (see figure 1.2, top). The Somalia plate diverges from both the Arabian and African plates, thus creating the Great Rift Valley, spreading in the Gulf of Aden through the Aden Ridge, creating an axis from the Tadjoura to Sheba Ridge (West to East). The Red Sea and Gulf of Aden are linked via the Bab-El-Mandeb Strait and are bordered by numerous mounts. The Eritrean and Ethiopian Highlands follow the western Chapter1. Presentation of the Northwestern Indian Ocean Chapter1. Presentation of the Northwestern Indian Ocean side of the Red Sea dominated by the Emba Soira (3018 m, the "mount of the priests") and Ras Dashen (4550 m, the "head guard") respectively. The eastern side is bordered by the Sarawat mountains, crossing Jordan, Saudi Arabia and Yemen and covered by the Jabal An Nabi Shu'ayb (the "mountain of the prophet Shu'ayb", 3666 m) near Sana'a. The Gulf of Aden is bordered South by the Galgodon Highlands (or Ogo mountains) that peak at 2416 m with the Shimbiri ("mount of the birds").

The Arabian plate and the Eurasian plate are in a convergent motion. This movement created the Zagros mountains, from the northern Iran down to the Strait of Hormuz, closes the Persian Gulf (see figure 1.2, bottom). The higher peak of this mountain range is the Zard Kuh (the "Yellow Mountain") with an altitude of 4548 m. Along the coast of Pakistan and Iran the Kirthar mountains link the Zagros range with the Himalayan Range. On the Oman coast this convergence created the Hajar mountains ("stone mountains") along its North coast, with the Jabal Akhdar ("the Green Mountain") as its peak.

The Indian Ocean has a mean depth of 3890 m but the northwestern part possesses three subsurface features along the frontiers of the tectonic plates. West, the Carlsberg Ridge is a divergent boundary between the Somalia and the Arabian plate that rises at 2000 m depth. North-West, the Owen Fracture Zone is a transform fault between the Arabian and Indian plates. North, the interaction between these two plates changes to a convergent zone. These fractures create shallowing up to 1000 m depth and deepening below 4000 m in the North of the basin.

Hydrological basins

Thus, the bathymetry and orography of the region define five hydrological basins : two semi closed basin, the Red Sea and Persian Gulf, linked to the Arabian Sea, via their respective gulfs, the Gulf of Aden and the Gulf of Oman (or Sea of Oman, used hereafter). During the following sections we divide these basins in three groups : the Red Sea and Gulf of Aden, the Persian Gulf and Sea of Oman, and the Arabian Sea. We focus now on their geophysical characteristics.

Red Sea and Gulf of Aden

The Red Sea, which name probably comes from the bloom of a red plankton, is oriented North-West to South-East. It extends from the Sinai Peninsula, that separates the Gulf of Suez and the Gulf of Aqaba, to the Strait of Bab-El-Mandeb, for a length of 2250 km and a width of 335 km. Along its axis, the Red Sea Rift digs to a maximum depth of 2211 m whereas the basin has a mean depth of 490 m.

The Bab-El-Mandeb Strait separates the Red Sea from the Gulf of Aden, with a minimal width of 25 km and maximal depth of 310 m.

The Gulf of Aden, deepens along its orientation from West to East. Maximum depths are found along the ridge from Tadjoura to Sheba, in the central part of the basin, with a peak at 2700 m. It also widens along its 1000 km length, from 150 to 440 km at its eastern part. This limit is defined in oceanography, as the continental shelf between the Socotra Island (12.5 o N; 54 o E) and the Kuria Muria Islands (17.5 o N; 56 o E), see figure 1.2 (top) for locations.

Climate

Persian Gulf and Sea of Oman

The Persian (or Arabian Gulf) has a length of 989 km from North-West to South-East and a width varying between 200 and 300 km. This Gulf is very shallow with a mean depth of 50 m and a maximum depth of 90 m along its northeastern, Iranian coast. On the East, the Persian Gulf opens on the Sea of Oman via the Strait of Hormuz. This Strait have a sill below 100 m depth in its southern part. Its narrowest width is of 39 km.

The Sea of Oman (or Gulf of Oman), as the Gulf of Aden deepens along its East-West axis to 3400 m in its western part. Also it widens from the Strait with 77 km, to 343 km at its mouth, for a length of about 500 km. This western frontier is formed by a line joining Ra's Al Hadd (22.5 o N; 61.75 o E) to Ra's Jiwani, at the frontier between Pakistan and Iran (see figure 1.2, bottom, for locations).

Arabian Sea

The Gulf of Aden and the Sea of Oman are the western and northern limits of the Arabian Sea, the northwestern region of the Indian Ocean. Its southern limit is defined by a line from Ra's Hafun, in the North-West of Somalia (10.5 o N; 51.25 o E) to the South of the Maldives (0.65 o S; 73.15 o E). Finally, its eastern frontier is limited by a line running through the Lakshadweep and Maldive archipelagos, between this limit and the Indian coast lies the Laccadive Sea. Its maximum depth is below 4600 m. The basin deepens sharply o↵coast, with the exception of the southern coast of Oman and along the eastern coast of India, where shallow plateau extends for 100 km o↵shore.

Climate

The strongest di↵erence between the Indian Ocean with the Pacific and Atlantic Oceans is the presence of a bordering continent in the northern hemisphere. This situation induces the creation of a ocean-atmosphere interaction, the monsoons. In this section, we present the e↵ects of this regime on the basins via the air-sea fluxes.

Arabian Sea

The Arabian Sea is strongly influenced by the monsoon that can be categorized in two distinct phenomena.

In summer, from June to September, the land temperature rises quicker than the ocean, creating an area of low pressure over the land. This di↵erence of pressure leads to sea breezes, winds from the sea to the continent. A northeastward jet develops during this period, parallel to the Somalia and southern Arabian Peninsula coast, called the Findlater Jet (from [START_REF] Findlater | A major low-level air current near the indian ocean during the northern summer[END_REF]). Summer monsoon brings steady winds from the South-West over the Arabian Sea with strong intensity, mean velocities being of the order of 14 m.s 1 . This monsoon is linked to the net heat loss in the central and eastern parts of the basin, with a maximum in June and July of -20 W.m 2 (see [START_REF] Montégut | Simulated seasonal and interannual variability of the mixed layer heat budget in the northern indian ocean[END_REF], [START_REF] Shaji | Seasonal variability of near-surface heat budget of selected oceanic areas in the north tropical indian ocean[END_REF] and [START_REF] Rao | Cooling and deepening of the mixed layer in the central arabian sea during monsoon-77 : Obersvations and simulations[END_REF]). In the western part, the net heat flux is positive (see Weller

Chapter1. Presentation of the Northwestern Indian Ocean et al. (1998)), this gain is principally due to the creation of upwellings along the coast of Somalia, Yemen and Oman (see [START_REF] Brown | Evolution of sea surface temperature in the somali basin during the southwest monsoon 1979[END_REF] and Fischer et al. (2002)). The southwestern winds create an Ekman transport, chasing waters, on their right, toward the South-East. This depression creates a vertical current, an upwelling, advecting deeper and colder water. These cold waters, steady winds and intense heating are the perfect conditions for strong evaporation. During the summer period, over 15 cm of evaporation is registered in June for example. A moist air mass rises, is advected by the winds and stopped by the Himalayan Range, leading to precipitation over India, from West to East during summer.

In winter, from November to February, the situation reverses, with higher pressure over the land, the winds blow steadily from the North-East with lower velocity (about 10 m.s 1 ). The air is dryer than in summer and brings very few precipitations over the Arabian Sea. Nevertheless it induces evaporation over the basin, as in summer, leading to a net loss of ocean heat. In the North-East part of the basin the evaporation is intensified by the trade winds from the North-East, bringing dry air over the ocean. This induces the formation of a surface water with strong salinity, called the Arabian Sea High Salinity Water.

In summary, both of these monsoons are dominated by strong, steady winds, net heat loss and evaporation of the basin, resulting in a loss of buoyancy of the sea surface (see [START_REF] Weller | Moored observations of upper-ocean response to the monsoons in the arabian sea during 1994-1955[END_REF]). The southern region also present the strongest winds, inducing a deepening of the mixed layer depth.

These two periods are separated by a transition regime with lower intensity and no clear wind orientation : inter monsoons. In spring, the transition of the North-East monsoon to the South-West monsoon takes progressively place, from West to East, with the weaker winds in the North of the Arabian Sea. In fall, this transition occurs rapidly, in only one month, October. The annual cycle of the sea surface temperature shows two phases of cooling, from mid-November to January and May to August, respectively linked to the lowest heating in winter and the presence of upwelled water in summer [START_REF] Shetye | Circulation and water masses of the arabian sea[END_REF]. From September to mid-November and February to March the sea surface water rises Due to the mountain ranges described earlier, the neighbouring basins are di↵erently impacted by the monsoons.

Gulf of Aden and Red Sea

The axis of the Gulf of Aden is lined up with the monsoon wind orientations and the surrounding mountains create a channel during the inter monsoon periods. Thus this basin presents winds with velocities over 5 m.s 1 all year long, from the South-West from June to August and from the North-East from October to April. The only exceptions are in September and May, associated with a reversal of the wind directions and the inter monsoon periods. The months of influence for the monsoons is slightly di↵erent from the Arabian Sea, due to the mountain ranges that influence the onset periods. This seasonal division is also available in the Red Sea under the influence of the Ethiopian and Eritrean Highlands on the West and the Sarawat Mountains on the East, thus creating a channel to the winds. North of the Red Sea (northern than 19 o N), the basin is not influenced by the monsoon regime and the winds blow from the North through the year (see [START_REF] Patzert | Wind-induced reversal in red sea circulation[END_REF]). South of 19 o N, winds are from the North in summer (June to 1.2. Climate August) and from the South-West, through the Gulf of Aden, in winter (October to April). These steady winds create conditions for evaporation over the Red Sea. The maximum values occur at the extremities of the basin, particularly in winter, with the exception of the reversal zone around 19 o N where evaporation is lower (see [START_REF] Neumann | Evaporation from the red sea[END_REF] and Privett (1959)). The annual mean was estimated at around 2 m.yr 1 . Furthermore the input of fresh water is low with no rivers flowing to it and low precipitations. The Red Sea exchange water only with the Gulf of Aden through Bab-El-Mandeb and a small amount via the Suez canal.

Sea of Oman and Persian Gulf

Being perpendicular to the monsoon wind directions and protected by the Zagros Mountain on the North and the Hajar Mountains on the South, the Sea of Oman is less subject to the monsoon than the other basins. Nervertheless, the presence of the Ra's Al Hadd cape (eastern cape of Oman) creates a rotation of the winds around it, thus being a region of intense variations of the wind stress curl. The winds rarely exceed 3 m.s 1 in the Sea of Oman, with northerlies from November to February, southerlies from June to September and westerlies during the inter monsoons in October and from March to May.

The Persian Gulf is located at the boundary between mid latitude and tropical climates, with descending dry air, thus creating arid conditions. All year long the Gulf is subject to winds from the North-East, the Shamal (meaning "North" in arabic). It blows numerous times between the end of spring and mid summer. In winter, the Shamal tends to appear abruptly, associated with high pressure above the Arabian Peninsula and produces strong winds (over 10 m.s 1 ) that last from three to five days (see Pous et al. (2013)).

Most of the evaporation occurs in winter due to strongest winds, but surface waters present higher temperature in summer. Privett (1959) estimates the evaporation to be on order of 1.44 m.yr 1 , which is coherent with the observations of Meshal and Hassan (1986). Reynolds (1993) points out that in winter the air is colder than the sea, thus creating a convective boundary layer. This unstable layer and the strong winds enhances the evaporation. In summer, the air is warmer than the sea and is humid; this buoyancy enhances evaporation during this season.

The climate being arid precipitations are very low, about 7 cm.y 1 . The rivers inflow, mainly Tigris, Euphrates and Karun via the delta of Chatt El-Arab and along the Iranian coast, varies from 35 to 110 km 3 .y 1 or equivalently to 15 to 46 cm.y 1 (Reynolds, 1993). But these values have to be renewed at the sight of the developing industry and agriculture in the region, demanding in fresh river waters. The most saline waters of the basin are created on the shallow coasts of the United Arab Emirates in the southern basin, an area with strong evaporation.

To summarize the atmospheric forcing provides steady, intense winds in the monsoon periods. These winds set upwellings, bringing fresher and colder water along the coasts of Somalia and Oman. Associated with these winds the net heat flux creates a loss of heat for the ocean via evaporation. In summer the moist air precipitates over the land, feeding the river input. On the adjacent, semi-closed basins, the Red Sea and the Persian Gulf, this evaporation greatly exceeds the input of fresh water through rivers and precipitations, thus generating saline waters.

Chapter1. Presentation of the Northwestern Indian Ocean

Oceanographic exploration, Historical discoveries and experiments

This region is and has been of high geopolitical interest, the Red Sea and Gulf of Aden being a direct way for exchanges between Europe and the Indian Ocean. Furthermore, the Persian Gulf and the neighbouring countries, as well as the Red Sea, are regions with the largest hydrocarbon resources. Adding to that, piracy along some territorial waters, and the ancient and recent conflicts from a local or international perspective, this whole region is of principal interest. But for the same arguments, having complete hydrographic experiments is rather di cult.

In the Red Sea the first observations were limited in time and mainly during the winter season in the North of the basin or only along the central axis (see [START_REF] Neumann | Evaporation from the red sea[END_REF] and [START_REF] Sofianos | Observations of the summer red sea circulation[END_REF]). The first recorded was from the Pola expedition in 1895 (from the Imperial Austrian Navy), followed much later by the Magnaghi expedition in 1923-1924 (from Italy) (from Italy), both with limited stations. In 1933In -1934, the John Murray/Mabahiss Expedition took place from the Arabian Sea to the Gulf of Aden and recorded several different water masses with a high number of stations (see [START_REF] Morcos | Oceanography of the gulf of aden: John murrays-mabahiss expedition 1933-1934 revisited[END_REF]). A basin wide campaign on the German ship Meteor took place in summer 1948 in the Persian Gulf, followed much later in winter 1976 by the Atlantis ship and Janurary to August 1993 for the Mt Michell expedition. From then on, numerous vessels took measurements in the 1950's such as the Discovery II and III (British), Norsel and Commandant Robert Giraud (French), A.I. Voyekov (ex S.S.S.R.) and Atlantis I, II and Vema (American) described by [START_REF] Neumann | Evaporation from the red sea[END_REF] among others. In 1962 to 1965, 27 nations took part in the International Indian Ocean Expedition, with the results studied in [START_REF] Wyrtki | Oceanographic atlas of the international indian ocean expedition[END_REF], setting in place the first atlas of this Ocean.

More recent experiments focused on regions that had not been sampled at specific periods. For example the R/V Maurice Ewing measured the Red Sea circulation in summer [START_REF] Sofianos | Observations of the summer red sea circulation[END_REF], or the PhysIndien 2011 measured circulation around the Arabian Peninsula during the spring intermonsoon. This experiment is described in this thesis. Some campaigns focused specifically on smaller regions but for longer time, such as in [START_REF] Murray | Direct observations of seasonal exchange through the bab-el-mandeb strait[END_REF] where they focused on the Bab-El-Mandeb Strait. The REDSOX (Red Sea Outflow Experiment) in 2001 had objectives to define the pathway and spreading of the Red Sea outflow.

Later, data assimilation in models, as well as the ARGO float program, allowed to have a better spatial coverage but some zones remain poorly sampled, such as the northeastern part of the Arabian Sea.

Water mass formation and circulation

We present here the circulation for each evaporation basin, and their exchange through the straits of Bab-El-Mandeb and Hormuz. Then we focus on the circulation in the Arabian Sea and how these water mass evolve.

1.4. Water mass formation and circulation

Red Sea and Gulf of Aden

Water masses formation

The geographical structure of the Red Sea, associated with the high insolation and with the steady winds, northerlies in the northern basin, and alternating southerlies and northerlies in the South, create conditions to strong evaporation leading to the formation of salty water. Two water masses with high salinity are formed in the basin, the Red Sea Deep Water formed near the Suez canal and in the Gulf of Aqaba (North of the basin); it occupies the bottom of the basin below 200 m [START_REF] Wyrtki | On the deep circulation of the red sea. Processus de Formation des Eaux Profondes[END_REF]; and the Red Sea Overflow Water in the northern part of the Red Sea, with a peak of formation in winter, and mainly due to oceanic convection [START_REF] Sofianos | An oceanic general circulation model (ogcm) investigation of the red sea circulation : 2. three-dimensional circulation in the red sea[END_REF]. These two waters show salinity over 40 psu and temperature around 20-22 o C (see 1.3, top, for ✓/S diagram in winter and summer). Two other water masses are found in the Red Sea, the Gulf of Aden Intermediate Water, with the lowest salinity and more present in summer as a subsurface layer. Finally the surface layer is occupied by the Red Sea Surface Water, saltier than the penetrating Gulf of Aden water, with temperature between 24-26 o C in winter and 28-30 o C in summer.

Circulation

In the Red Sea, the horizontal circulation is dominated by large eddies, about the width of the basin with maximum velocities of about 1 m.s 1 . Their formation is associated to the strong wind stress curl [START_REF] Quadfasel | Gyre-scale circulation cells in the red sea[END_REF]. These eddies are mostly anticyclonic in the center of the basin, but on its northern side, a cyclonic gyre remains stationary all year long [START_REF] Cli↵ord | An oceanographic nowcast/forecast system for the red sea[END_REF] and participates largely in the formation of the Red Sea Overflow Water [START_REF] Sofianos | An oceanic general circulation model (ogcm) investigation of the red sea circulation : 2. three-dimensional circulation in the red sea[END_REF]. This water mass is advected southward around the eddies. In contrast, the inflowing Gulf of Aden Intermediate Water extends to 16 o N, up to where the gyres start to be the main features, then it is subject to strong erosion. Coastal currents are also observed but only seldom.

The southern part opens on the Strait of Bab-El-Mandeb, with a sill at 160 m depth. Its circulation is influenced by the seasonal reversal of the monsoon with a maximum of Red Sea Outflow Water in winter (about 0.6 Sv; [START_REF] Patzert | Wind-induced reversal in red sea circulation[END_REF]) and minimal in summer. In winter (October to April), the exchange through the strait is composed of two layers, the surface layer being occupied by inflowing water from the Gulf of Aden, fresher and colder, and an underlying, outflowing, layer of Red Sea Water, saltier (larger than 40 psu) and warmer (see Al Saafani and Shenoi (2004)). In summer, the northerly winds create a thin third layer above salty and warm outflowing Red Sea Surface Water with salinity of 36.5-37 psu and temperature up to 30 o C in the Gulf of Aden (see figure 1.3, bottom).

In the Gulf of Aden, the Red Sea Outflow Water plunges and divides in two branches, a first near Djibouti sinks in the Tadjoura Rift (Bower et al., 2005) and a second follows the coast of Yemen, up North (Bower et al., 2000). The water mass equilibrates between 600 and 800 m with an intermediate layer salinity maximum at 650 m (Bower et al., 2005) around a potential density ✓ of 27 to 27.4. The horizontal circulation is also dominated by mesoscale eddies with a width comparable to the size of the basin (about 200 km in diamter) and a strong vertical influence (see Bower et al. (2002)). They present a maximum velocity of ⇡ 0.4 m.s 1 and a westward propagation. Their formation sites are supposed to be in the western Arabian Sea, related with the wind stress curl. Due to their barotropic signature they have a role on the advection of the Red Sea Outflow Water out of the Gulf of Aden in the Indian Ocean. In summer, the presence of the cold and Chapter1. Presentation of the Northwestern Indian Ocean 1.4. Water mass formation and circulation fresh coastal upwelling water creates a front with the saltier and warmer Red Sea Outflow Water at the mouth of the Gulf of Aden.

Numerical Studies

To study the outflow from the Red Sea and the Persian Gulf (Bower et al., 2000) used an outflow model PB94 that described the outflow as a density driven single layer penetrating an open ocean stratified and at rest. It focused on the impact of seasonal variation of water source or oceanic properties, topography and low latitude e↵ects on the resulting outflow. The results for the Persian Gulf outflow are described at the end of the next section. In the Gulf of Aden, two di↵erent topographic e↵ects are underlined. First, the channel along the coast of Yemen, with a slight slope produces the deep layer of Red Sea Water, between 600 and 800 m. Second, the central sill produces the water on the surface layer, down to 500 m. The bottom layer, principally in the Tadjura Rift, were not studied in this model.

More recently, a MITgcm (general circulation model) simulation was laucnhed with the aim of studying the circulation and season and inter annual variability in the Red Sea. This simulation have a horizontal resolution of 1.8 km, 25 isopycnic vertical level, high frequency atmospheric forcings (NCEP reanalysis with a 6 hours recurrence) and realistic boundary conditions. The seasonality and characteristics of intrusion of Gulf of Aden Intermediate Waters in the Red Sea from the model are regrouped in [START_REF] Yao | Seasonal overturning circulation in the red sea: 1. model validation and summer circulation[END_REF] for summer and Yao et al. (2014a) for winter.

In summer, the colder and fresher water from the Gulf of Aden inflows as an intermediate layer along the eastern coast, being advected by mesoscale eddies in the northern basin. These waters are trapped inside the structure and mixed with the surrounding waters during the reversal of the monsoon. The forcing that drives the penetration of these waters is the shallowing of the thermocline in the Gulf of Aden, more drived by the monsoonal winds above the Indian Ocean than the upwelled waters. In the Red Sea, the penetrating winds enhance the outflowing surface layer of Red Sea high saline waters. This induces a slope of the surface layer northward, thus creating a baroclinic pressure gradient advecting inflowing, less saline water as an intermediate layer. The study of the heating di↵erence between Gulf of Aden and Red Sea reveals a change of the surface temperature and salinity in both of the basin in late summer, modifying therefore their vertical structures.

The results for the winter simulation in Yao et al. (2014a) point out the presence of a overturning circulation in the northern part of the basin, composed of upwelling along the western coast and downwelling on the East. Gyres forms in the boundary layer of 20 km and play a role in the increasing density of the sinking waters.

Persian Gulf and Sea of Oman

The Persian Gulf is very shallow thus making it subject to strong tides (Pous et al., 2012). Tides have a strong influence on small horizontal and time scales (<10 km and <24 hours). They contribute to vertical and horizontal mixing and stirring, with elevation higher than 1 m in all the basin, but on a temporal average they have very little influence on the basin scale circulation. The Persian Gulf has two diurnal amphidromic points, one in the North-West of the basin, the second in the South between Qatar and the United Arab Emirates, and one semi-diurnal amphidromic point in the center, North of Bahrain. These In winter the northern Persian Gulf is dominated by vertical mixing, with the exception of the river plume from the Shatt El-Arab. Vertical mixing is mainly influence by the tides, the evaporation and the winds and bottom frictions. The tides create a bottom mixed layer whereas evaporation and wind e↵ects have a stronger influence on the surface layer. In winter, the strong evaporation associated with surface cooling creates high density water that sinks and mixes the entire water column. The di↵erence in density over the basin is principally due to the di↵erence in salinity between the saline water of the Persian Gulf and the fresher water from the Sea of Oman whereas the di↵erence in temperature contributes much less.

In summer the stratification of the Persian Gulf is stronger and a thermocline can be found around 18 m in depth, excepted in the northern part. Below and above this thermocline, well mixed waters can be found, the densest, at the bottom, being formed in winter. Near the Shatt El-Arab delta, river plumes with low salinity concentrate on the surface layer. The dense water flows southward, between the thermocline and the bottom. In summer the di↵erence of density across the basin is sustained by both temperature and salinity, contrary to winter where principally salinity plays a role.

In summary, the northern part of the Persian Gulf, shallower, is more influenced by the wind e↵ects and to the river run o↵s, whereas the Southern part responds to the density e↵ects and the reverse estuarine flow via the Strait of Hormuz. The Gulf provides water with strong thermohaline characteristics, water with salinity over 40.5 psu in winter and water with salinity around 40 psu but warmer in summer (see 1.4) In the Strait of Hormuz, between the Persian Gulf and the Sea of Oman, the dense water from the Persian Gulf is stratified in winter and flows out at the bottom with a temperature of 20 o C and a salinity of 40 psu. Near the surface, along the Iranian coast, the water is not stratified, but the isopycnals correspond to a geostrophic current towards the North. In summer the outflow across the Strait of Hormuz is stratified; the tides create a mixed bottom layer with a temperature of 20 o C and salinity of 39.5 psu. The strong tides within the strait mainly contribute to decrease the stratification and create strong vertical mixing (see Banse (1997)). This reverse estuary flow is comparable to the circulation between the Atlantic Ocean and the Mediterranean Sea via the Strait of Gibraltar. Other forcings are due to the fresher water from the rivers flowing southward; along Iran, a coastal current is advected northward.

Once in the Sea of Oman the Persian Gulf Water sinks and equilibrates below 200 m in depth with temperature and salinity that peak at 21 o C and 37.5 psu with a thickness of 100 m. Banse (1997) compared the core of the Persian Gulf outflow from several expeditions and associated it with a range of density ✓ varying from 26.30 to 26.95. In Banse (1997) it was suggested that due to inadequate spatial and temporal coverage, the spreading of the Persian Gulf Water in the Sea of Oman, and furthermore in the northwestern Arabian Sea, could only be explained in very general terms. Reynolds (1993) described that the Persian Gulf Water outflow follows the coast of Oman and sinks along the shelf to form a layer of high salinity water near 260-270 m. The circulation in the Sea of Oman is widely studied in the following chapters.

Chapter1. Presentation of the Northwestern Indian Ocean

Numerical Studies

The Persian Gulf outflow in the Sea of Oman is also studied in the PB94 model used by Bower et al. (2000). The summer outflow is warmer than winter of ⇡ 3 o C and 0.2 psu fresher, but in the Sea of Oman they both have the same temperature, mainly dominated by the oceanic temperature from the Indian Ocean Central Water. In summer, the density of the Persian Gulf outflow is ✓ = 26.6 whereas in winter it is ✓ = 26.3, this di↵erence being due to the saltier Persian Gulf Water in summer in the Sea of Oman (greater by 0.5 psu). Nevertheless, due to the stratification of the surrounding waters, the outflow equilibrates around 250 m depth.

Using HYCOM, a hydrostatic, primitive equations resolving model, Yao and Johns (2010a) studied the Persian Gulf and the Sea of oman with a horizontal resolution of 1/20 o and a vertical resolution in the z coordinate for the weak stratification layer and in the isopycnal coordinates below. Yao and Johns (2010a) use di↵erent atmospheric forcing to study the pattern of the seasonal circulation and the water mass formation. A first statement is the propagation of Indian Ocean Central Water that dominates the surface circulation in the Persian Gulf, within two branches along southern coast of the basin and the Iranian coast. The first branch veers under the influence of the Ekman drift, and the second is reinforced in summer. This Indian Ocean Central Water inflow forms cyclonic gyres in the northern and southern part of the basin. A second statement is the formation of strong salinity water from the loss of heat and fresh water, principally on the shallowest part of the Persian Gulf, in the northern and southwestern part. The simulation also provides a description of the exchanges through the Strait of Hormuz (see [START_REF] Yao | A hycom modeling study of the persian gulf : 2. formation and export of persian gulf water[END_REF]). In winter, the exchange is as described earlier in the Mt Michell experiment (see Reynolds (1993)), a two-layer exchange, with the dense salty water in the bottom and along the southern part of the Strait and a surface inflow, planked North, of Indian Ocean Central Water. Nevertheless, in summer, an outflow at intermediate depth of salty and warm waters formed in the southwestern part of the Persian Gulf is observed.

Also using HYCOM but focusing on the Strait of Hormuz with a 1 km resolution and 16 designed vertical level, Thoppil and Hogan (2009) focused on the Persian Gulf outflow. A wind forcing with a high resolution is used as well as a GDEM climatology for thermohaline characteristics. It results in the observations of pulses of high salinity outflow principally in spring and summer with a period of 15-30 days and an amplitude of 2 psu. These peaks are linked with the fluctuations of the wind stress, thus creating variations of the flow around Hormuz. These fluctuations induces local instabilities forming mesoscale cyclonic eddies with a diameter of about 65 km that are coherent with barotropic instability. Along the Strait they advect heat and salt into the Sea of Oman before being eroded by vertical mixing.

Using the same configuration but focusing on the Persian Gulf [START_REF] Thoppil | A moedeling study of circulation and eddies in the persian gulf[END_REF] studied the mesoscale circulation of the basin. In summer a cyclonic circulation takes place, extending northward until July. At the same period the Iranian coastal current is at its maximum in width and velocity. The penetrating fresher and warmer Indian Ocean Central Water creates a density front that destabilizes the current, thus deforming in meanders and mesoscale eddies. These eddies are cyclonic with size of 110 km, in agreement with a baroclinic instability. In July and August these eddies are observed along the coast of Iran before dissipating in winter.

Water mass formation and circulation

Arabian Sea

Arabian Sea Circulation

The Arabian Sea is influenced by the monsoon regime. The seasonal changes of monsoon influence the coastal currents and the setting of upwellings. These latter occur in summer, principally along the southern coast of Oman and along Somalia and uplift colder and fresher water.

During winter, the circulation is globally cyclonic. The North Equatorial Current flows westward between the Equator and 10 o N and brings water with low salinity from the Bay of Bengal into the Arabian Sea. A branch of this current turns South of India (Shetye et al., 1991) to form the West India Coastal Current around the Lakshadweep High, an anticyclonic eddy. On the western side of the basin, the winds force a southward coastal current along Somalia to the Equator and along the coast of Yemen and Oman, the Oman Coastal Current flows southwestward. To close this circulation, the Counter Equatorial Current flows eastward South of the Equator.

In summer the winds are intense along the Somalian coast and form the Findlater jet, blowing from the South-West, deepening the mixed layer and forming upwellings. In the same region, two anticyclonic eddies appear, the Great Whirl and the Socotra Eddy. The North Equatorial Current is replaced by a stronger Indian Monsoon Current flowing eastward. Along the South of India, a cyclonic eddy, the Lakshadweep Low, sets during summer, and the West Indian Coastal Current flows southward. Along the coast of Yemen and Oman, an upwelling rises colder water, advected northeastward under the influence of the Oman coastal current.

The cycle of the Lakshadweep, with cyclone and anticyclone, is related to the coastal Kelvin waves from the Bay of Bengal that radiate into Rossby waves o↵ southern and western India (see [START_REF] Shankar | Hydrography of the eastern arabian sea during summer monsoon 2002[END_REF]).

In summary, the basin scale circulation reverse from winter to summer, from cyclonical to anticyclonical, and is dominated by three stronf coastal currents that induces the formation of gyres. The Somalia Coastal Current forms the Great Whirl in summer near Cape Guardafui. The West Indian Coastal Current near Cape Guardafui induces the formation of the Lakshadweep high and low, in summer and winter, near Cape Kanyakumari, or Cape Comorin. Along the coast of Oman and Yemen, the Oman Coastal Current is southwestward in winter and northeastward in summer. Near the Ra's Al Hadd Cape, the western point of the Arabian Peninsula, the Oman Coastal Current ejects eastward in summer under the influence of the Ra's Al Hadd jet, a dipole composed of a cyclone North of the cape, and an anticyclone South of it. The coastal circulation along the northern Arabian Sea, along the coasts of Iran and Pakistan is poorly described. Finally, the Equatorial Current is westward all year long, but in winter the Counter Equatorial Current takes place North of it, with a motion towards the East.

High Salinity water masses in the Arabian Sea

The previously described Red Sea Outflow Water and Persian Gulf Water flows in the Arabian Sea all year long. We describe hereafter the characteristics of the di↵erent water masses that show high salinity in the basin. Three peaks are observed : near the surface, the Arabian Sea High-Salinity Water, between 200 and 300 m, the Persian Gulf Water and between 500 and 800 m the Red Sea Water (see figure 1.5). These waters dilute away from their basin of formation and unite as a thick layer between 150 and 900 m of North

Water mass formation and circulation

Indian High-Salinity Intermediate Water.

On an annual mean, Prasad et al. (2001) found that Persian Gulf Water penetrates between 200 and 250 m, with a salinity decreasing from 36.7 to 35.1 psu from the Sea of Oman towards the equator. Similarly, the temperature goes from 17 o C to 13 o C southward. This water mass propagates on ✓ = 26.2-26.6. Prasad et al. (2001) also suggests that the standard deviation of the Persian Gulf Water depth is smaller than 40 m in all the Arabian Sea, with the exception of the southwestern part (near 10 o N; 60 o E) where the strong negative wind stress curl, associated with the Findlater jet, deepens the water mass below 300 m. The lowest thermohaline properties of the Persian Gulf Water are found near the western Indian coast, where the mixing with the fresh Bay of Bengal waters keeps these characteristics low, particularly in winter. The pathway of the Persian Gulf Water is modified by the season. In winter, it exits South the coast of Oman then spreads eastward. During the spring inter monsoon, the core of Persian Gulf Water deepens from 240 to 280 m. During the summer, the core of ✓ = 26.5 decreases to reach 200 meters near the upwelling regions, along the Coast of Yemen and Somalia. These regions also present lower salinity due to vertical mixing and fresher water advected northward along the coast of Somalia, around the Great Whirl. In contrast, the observed Persian Gulf Water in the center of the Arabian Sea deepens as the mixed layer depth, with an anticyclonic motion. In autumn, the Persian Gulf Water is advected out of the Sea of Oman to the South, o↵ Ra's Al Hadd. Then it spreads along two paths, a first one towards the South-West along the coast of Oman and a second one towards the South-East.

The Red Sea water penetrates the Arabian Sea through the Gulf of Aden with a temperature of 11 o C and 35. , between 600 and 800 m. The circulation of this water mass is influenced by the monsoon with a propagation southward along the coast of Somalia and in summer rather eastward.As the Persian Gulf Water, Red Sea Outflow layer in the central Arabian Sea deepens during the monsoons, under the influence of the strong winds.

Both the Red Sea Outflow Water and the Persian Gulf Water have southward motion and their spreading is not influenced by the variability in their thermohaline characteristics (see [START_REF] Beal | Spreading of the red sea overflow waters in the indian ocean[END_REF] and Prasad et al. (2001)).

The Arabian Sea High Salinity Water forms directly in the Arabian Sea, oppositely to other water masses with strong thermohaline characteristics, but whith di↵erences through the seasons. This induces a lighter density, in summer, ✓ = 23.5-23.9, associated with a temperature of 26-27 o C and a salinity of 36.5, observed as the lightest water on figure 1. 5 (top and bottom). In winter this water mass is heavier with colder water from 23 to 25 o C, a salinity of 36.6 psu and theta = 24.3-24.8. This water mass is also more influenced by the seasonality of the region. The formation of the Arabian Sea High Salinity Water occurs in winter and extends towards the South, deepening from the surface to 60 m at 13 o N while mixing with waters with lower salinity (see Kumar and Prasad (1999)). Still in early winter, a second patch can be found near 13 o N; 68 o E, a remnant of the previous year. In spring, the core deepens everywhere at 60 m and stays around 65 o E. The northward current along the coast of India prevents the high salinity water to spread East. In early summer, the water shows little extent, in the North-East of the basin. Then it spreads southward until earlier winter. The cold and low salinity upwelled waters from the South-West monsoon create a front with the Arabian Sea High Salinity Water (see [START_REF] Montégut | Role of fronts in the formation of arabian sea barrier layers during summer monsoon[END_REF]). Furthermore, the northern part of the Findlater Jet, associated with a positive wind stress curl shoals the core of the high salinity water in the northern Chapter1. Presentation of the Northwestern Indian Ocean part of the basin.

Low Salinity water masses in the Arabian Sea

The moist air over the Himalayan Range precipitate mainly along the eastern coast of India, in the Bay of Bengal, about 2 m.yr 1 [START_REF] Prasad | Annual and seasonal mean buoyancy fluxes for the tropical indian ocean[END_REF], and by rivers run o↵. The Bay of Bengal receives the outflow from the Ganga and Brahmaputra present an outflow of 1300 10 9 m 3 .yr 1 [START_REF] Sengupta | Surface freshwater from bay of bengal runo↵ and indonesian throughflow in the tropical indian ocean[END_REF], and the total river run o↵ is of 2950 10 9 m 3 .yr 1 . This input is over 60% of the freshwater influx in the northern Indian Ocean. Along the western coast of India, the maximum river input is from the Indus (80 10 9 m 3 .yr 1 , followed by the Gulf of Khambhat that receive waters from the Narmada, Tapi, Mahi Sabarmat and Ozat rivers, respectively, 32, 11, 7 and ⇡ 1 10 9 m 3 .yr 1 . These river fluxes are calculated from http://www.sage.wisc.edu/riverdata/index.php?qual=32 . Their maximum outflow are reached during the summer monsoon, linked with the strong precipitations.

From the eastern side of India, these waters contributes to the formation of the Bay of Bengal Low Salinity Water, a surface layer with mean thermohaline characteristics of 25-29 o C, 28-35 psu and ✓ = 21-22. This very light water mass is advected South by the East Indian Costal Current in winter and penetrates the Arabian Sea up to 16 o N under the influence of the West India Coastal Current (see Prasanna [START_REF] Kumar | Intrusion of the bay of bengal water into the arabian sea during winter monsoon and associated chemical and biological response[END_REF] and [START_REF] Rao | Characteristics of bay of bengal water mass in south eastern arabian sea during 2001-2002[END_REF]). It gains salt through mixing and is present in the eastern Arabian Sea with a density ✓ = 22.5-24.

From the Gulf of Khambhat, on the West coast of India, the maximum of outflow is also observed during the summer monsoon, but the southwesterly winds from this seasons keep this water mass inside the gulf. This does not apply to the Indus river that spread its fresh water southward along the coast of India during this season. Post-summer monsoon, the overturning of the wind direction and of the West Indian Coastal Current advects these waters North (see [START_REF] Rao | Observed low-salinity plume o↵ gulf of khambhat, india, during post-monsoon period[END_REF]).

The water mass that presents the lowest T/S characteristics and occupates the main core of the Arabian Sea is the Indian Ocean Central Water (see Tomczak and Godfrey (2003). Schott and McCreary (2001) stated that this water mass forms in the Subtropical convergence zone, migrates westward under the influence of the Equatorial Current and penetrates in the Arabian Sea with the Somalia Coastal Current. It extends from 100 m to below 1000 m depth. It presents a wide range of density with t heta from 25.5 to 27.

The bottom layer of the Arabian Sea to 1500-2000 m is occupied by Antarctic Bottom Water from the Circumpolar Currents and the Indian Deep Water, a mixing between the North Atlantic Deep Water and the Antarctic Bottom Water. These water penetrates the Arabian Sea from the South-West and are not clearly distinct, from their same circulations in the basin and their mixing characteristics (see [START_REF] Morales | Variability of water masses in the western indian ocean[END_REF]

The Phys-Indien 2011 Experiment

The Phys-Indien 2011 program, directed by the French Hydrographic and Oceanographic Service of the Navy (SHOM), was designed to measure the circulation and the thermocline properties of the di↵erent water masses around the Arabian peninsula. Between February and March 2011, hydrological and current profiles with Vessel Mounted ADCP and Sea-Soar were performed in the Red Sea, the Gulf of Aden, the Arabian Sea and the Sea of Oman and floats were deployed in the main water masses. We present hereafter the data collection of the program, characteristics and processing, as well as the di↵erent databases we used for our studies.

SeaSoar and Probes

Devices specifications

In the Red Sea and the Gulf of Aden, long transects across the basins registered temperature and conductivity with Expendable Bathy-Thermograph (XBT) and Expendable CTD (or XCTD), and velocity via VM-ADCP. In the Arabian Sea, sections were performed with the SeaSoar and VM-ADCP both along and perpendicular to the coast of Oman. The SeaSoar measures the thermohaline characterics of the water with two sensors (SBE 21 SeaBirds), both with an accuracy of 0.01 o C in temperature and 10 3 S.m 1 in conductivity. In order to calibrate the data of the SeaSoar, probes were cast with di↵erent resolutions and accuracy. The CTD SBE911 measure conductivity, temperature and pressure with the higher accuracy (10 3 o C, 3 10 4 S.m 1 and 0.015% of the pressure). The XCTD and the XBT have a lower accuracy but also a lower cost. The accuracy of the XCTD is 0.02 o C ans 3 10 3 S.m 1 for two di↵erent models, XCTD 1 and 2 that sink at 1000 m and 1850 m respectively. The XBT has an accuracy of 0.1 o C and three types were used XBT T5 T7 and T10 that dives at 1830 m, 760 m and 200 m depth respectively with vertical resolution of 65 cm and 2 % accuracy of total pressure for XCTD and XBT.

Chapter2. Data collection and Processing

Data Processing

The SeaSoar measured temperatures and salinities along the sections with a measure every 3 s from 50 m to 360 m of the water column with successive rising and lowering. A visual analysis of data showed crenelations in some of the salinity profiles (see SeaSoar section on figure 2.2, up). These errors arise from three separate sources (see [START_REF] Larson | Oceanographic ctd sensors : Principles of operation, sources of error, and methods for correcting data[END_REF] and Barth et al. (1996)), a problem of alignment between the temperature and conductivity measurements, a di↵erence of transfer of heat in the captors between the lowering and rising of the device, and a mismatch of the time constant of the measure- Chapter2. Data collection and Processing ments of temperature or conductivity. An automatic processing from ODV do not remove those errors and give too important smoothing, compared to the CTD, principally due to the di↵erent water masses from the basin. Therefore a correction of the salinity based on the measurement of pressure, temperature and conductivity is applied using the Barth et al. (1996) method. The first problem of the alignment between the data can be easily corrected with an o↵set. A couple of coe cients are used to correct those crenelations, one for the amplitude of the measurements between respective rising and lowering, and a second one for the time delay between temperature and conductivity (see Lueck and Picklo (1990)). For each section, in order to find the best values for these coe cients, we minimize the temperature and conductivity di↵erences between a lowering and a rising of the SeaSoar.

To validate this processing, cross correlation of the new values with CTD casts at the same locations are processed. The data is then interpolated on the depth to obtain 2-D matrices and filtered to suppress local density inversions (example on figure 2.3). Biases can still be observed due to the gap in location and time of the measurements of the CTD and SeaSoar, as well as the interpolation and filter that take into account neighbouring values. This processing allowed the correction of the crenelations without smoothing or loosing the peaks (see figure 2.2, bottom).

VM-ADCP

Device specifications

The Vessel-Mounted ADCP provides high resolution current profiles on two frequencies, 38 kHz and 150 kHz in the upper ocean (respectively in the 0 -1000 or 1400 m depth range, and in the 0 -250 or 350 m depth range) with a 1 % accuracy or ± 5 10 3 m.s 1 accuracy. A first analysis of the raw data reveals errors such as missing values on several meters of the water column, or excessive velocities at some points. Furthermore, a few profiles do not allow an automatic processing. Also, for some hours of the night, profiles below 800 meters are totally blanked due to biological activities known as the diel vertical migration (see figure 2.4).

Data Processing

As for the salinity and temperature from the SeaSoar, an automatic processing from ODV smoothed too much the data with final maximal velocities of 40 cm.s 1 where the double was measured. A first step in the data processing is to remove the values showing di↵erences with the neighbouring measurements exceeding five times the mean standard deviation. This threshold is high enough to safely keep all the signal and remove only the spurious measurements. After the erroneous values are removed, a median filter is applied to fill the blanks. This filter takes into account 4 points, above below, right and left, and weighted their values by the distance from them to the blank. (see figure 2.5, up).

The second step is to filter out the noises from the signal that can be due to very fine scale dynamics, biases in the measures or biological activities. Then we process the spectral signal in order to reduce the noise and keep the submesoscale to larger scale structures, scales above 10 km. We developed a low pass filter with some focus, the time step between each recording of the VM-ADCP showed variations and the direction of the ship also varied during the sections. These track interruptions are indicated by thick lines on the figure of the VM-ADCP. We apply a low pass filter function f = exp ( ( k k 0 ) 10 ) with k the wavelength and k 0 the cuto↵ wavelength. We used k 0 = 2⇡ 30 10 3 . An example can be seen on figure 2.5 (bottom).

Floats

During Phys-Indien, five drogued Surdrift buoys were deployed and programmed for 180 days of recording. A large holey-sock drogue was fastened to the surface buoy by a thin Kevlar cable, 80 to 250 m long. A test on the relative acceleration of each buoy was applied to determine a possible loss of the drogue. They were positioned by Argos, their trajectories were sampled every hour and a thermistor sensor gave the surface temperature at each recording.

As well, 6 PROVOR floats were deployed during this campaign. They were positioned via Argos. A CTD probe measured temperature, conductivity and pressure at each recording, with an accuracy of 0.01 o C, 10 3 S.m 1 and 1 dbar respectively. PROVOR are typically designed to have a cycle of 10 days, for 5 years. Here, they were re-calibrated to a 5 day cycle. Their parking depth were also programmed to drift at 700 dbar instead of 1000 dbar; then they dived again to 2000 dbar and finally surfaced to transmit their data via satellite. Chapter2. Data collection and Processing

Climatologies and long time observations

The Phys-Indien provided data for the spring 2011 period, nevertheless, climatological thermohaline characteristics and long time period measurements from floats and satellites are also necessary to compare those shorter period data. Hereafter is a description of these longer time data.

ARGO floats and Climatology

We also obtained ARGO float data from the Coriolis data center of IFREMER (France) with the ANDRO database. They provid temperature et salinity sections from the surface to 2000 m deep along their trajectories. Their lifetime range from a few months to several years. The floats from IFREMER have been processed (see Ollitrault and Rannou (2013)), but some were retrieved from earlier period and required a few filtering and control of the profiles. The control method used is the same as the VM-ADCP, without the low-pass filter. The number of float increased progressively during this thesis. In December 2014, we used all the floats from the ANDRO database that sampled the Arabian Sea, Gulf of Aden or Sea of Oman, about 300 floats, with periods varying between 5 and 10 days, between 2002 and 2014.

We retrieved the Generalized Digital Environment Model (GDEM) climatology from http://www.usgodae.org/ and used it to create horizontal maps of temperature and salinity with 1/4 ⇥ 1/4 degree resolution and a good coverage of the Sea of Oman or the Gulf of Aden. The accuracy of the GDEM climatology is of 0.5 o C in temperature and 0.1 psu in salinity for the 400 first meters then 0.25 o C in temperature and 0.05 psu in salinity until 2000 m [START_REF] Teague | A comparison between the generalized digital environmental model and levitus climatologies[END_REF]). These maps are then modified by inserting the float data. First, each float measurement is converted into an anomaly by substracting the climatological value for the given time and location. Second, this pointwise anomaly is extended spatially with an isotropic Gaussian correlation function, weighted by the local number of data (the radius of the Gaussian was chosen as L = 50 km, typical of a mesoscale eddy). This correlation function is normalized to unity where data lay. Third, the resulting anomaly was added to the climatology (see figure 2.6). Those maps are useful to study the variations with a mesoscale resolution and must be compared to maps of representativity (see figure 2.7) that show the number of float and their distributions. These maps put in focus some regions that are poorly sampled during experiments and are not integrated in the GDEM climatology, as the coastal features or strong local thermohaline gradients. As an example, the southwestern part of figure 2.6 has few measurements of the ARGO floats, it tends to represent the climatology from GDEM, but the transition between those two field is not physically accurate. On the contrary, the Sea of Oman has a good floats coverage and represent a westward gradient of temperature and salinity along the coast of Iran that was not described in the GDEM climatology.

Altimetry

Satellite description

The altimetric maps are obtained from the AVISO satellite data center with a 1/4 ⇥ 1/4 degree resolution on a Mercator grid and a daily interpolation. The mean sea level was a 22-year mean of the sea surface topography. The error on along-track measurements of sea level was 3-4 cm (Fu and Cazenave (2000)). Three altimetric satellites ( and 2, Envisat) are used. For the Jason satellites, the inter-track distance at the Equator is 315 km and the period of the orbit is 10 days to record the same track; for Envisat, they are 80 km and 35 days. From the along-track data, gridded data were computed by AVISO via data merger and optimal interpolation (see Ducet et al. (2000)). The quality of the maps strongly depends on the quality of the mean topography of the ocean surface, calculated from hydrographic and satellite data. More information on the Mean Absolute Dynamic Topography (MADT) production is provided in Gaspar and Ogor (1996), [START_REF] Traon | Mesoscale mapping capabilities of multiple-satellite altimeter missions[END_REF], Le Traon et al. (2001) and Labroue et al. (2006).

Processing

To identify eddy structures in altimetry , we use maps of MADT anomaly, which is the di↵erence between local MADT and its instantaneous spatial mean. This is di↵erent from the Sea Level Anomaly where the pointwise time average is retrieved from the MADT (see figures 2.8 and 2.9). We also use a statistical point of view using the Empirical Orthogonal Function (EOF), that decompose a time series of measurements on a 2-D field into its principal modes of variability. This provides us the spatial structure of variability. Geostrophic surface currents are computed from the MADT anomaly; such as the deep currents with hydrology from modified GDEM, deformation maps and Okubo-Weiss quantity maps.

Chapter2. Data collection and Processing

Numerical modelling

To complement this thesis and to focus on finer scales we used numerical simulations performed with the HYCOM model. Runs were performed by the SHOM in Toulouse. HYCOM is a primitive equation model with five equations, two for the horizontal velocity components, one for the continuity and two for the conservations of salt and temperature (see [START_REF] Bleck | An oceanic general circulation model framed in hybrid isopycnis-cartesian coordinates[END_REF] and Chassignet et al. (2007)). "HYCOM" stands for hybrid coordinate ocean model : Hybrid because levels with low stratification are represented with z coordinates and deeper levels are isopycnic or topographically stretched. The model has 40 vertical levels. The horizontal resolution is here 5 km from 5 o N to 30 o N and 32.3 o E to 77 o E, thus covering all the basin of interest here. A simulation of 11 years was produced, starting in 2011.

The model variables are discretized on an Arakawa C-grid with the mass related quantities evaluated at the center of the grid and the velocities on the faces of the grid.

The simulation was initialized with a state given by a lower resolution realist operational global model (MERCATOR). The southern and eastern boundary conditions were forced by this climatology. The atmospheric forcing (winds and heat) were obtain from the realistic fields from the year 2011 (from the French Met'O ce METEO FRANCE). They have of 1/8 ⇥ 1/8 horizontal resolution and 6 hours temporal resolution and a filter was applied to delete discontinuities between the end and beginning of year 2011. The real tides from 2011 to 2021 are forced as boundary conditions.

From the 11 years simulation, 5 years correspond to the spin-up of the model and are not used hereafter. The spin-up is over when the di↵erent water masses of the basins, described in the presentation chapter, have values correlated with the climatology, in thermohaline characteristics, locations and depth; as well the spin-up is achieved when the mesoscale circulation is observed and the peak of kinetic energy becomes more constant with seasonal variations. Only the last six years, 2016 to 2021 are analyzed.

Comparisons between the model results and the observations are made in chapter 5 of this thesis, with a focus on the Persian Gulf outflow and the mesoscale circulation in the Sea of Oman and northern Arabian Sea.

Chapter2. Data collection and Processing

The ROMS simulation and a spectral model of the 2D flow equations, were also used to study processes, particularly in chapter 5 and in the appendix on the Gulf of Cadiz.

Motivations for the research

The surface circulation in the Arabian Sea and Gulf of Aden has been described as dominated by mesoscale gyres (see Bower et al. (2002), Fratantoni et al. (2006) and Al Saafani et al. (2007)) with a strong seasonality, under the influence of the wind regime. As well, the outflowing waters from the Red Sea and Persian Gulf have noticeable thermohaline characteristics; these characteristics and spreading are influenced by the monsoons. An aim of this thesis is to study their variability, on a mesoscale level, of those gyres, and their influence on the water masses.

Chapter3. Mesoscale variability of water masses in the Arabian Sea as revealed by Argo floats

Observations over long periods and on regional scale scale are needed. An Eulerian study of the our basins of interest over several seasons is di cult to achieve. As stated in the presentation of the region, experiments are relatively rare and are focused on specific region or seasons. Therefore, we choose to use a Lagrangian approach by using ARGO floats. This allows us to have first a qualitative approach of the mesoscale circulation, by focusing on few floats.

The circulation in the Gulf of Aden being well described, we focused our first study on this region, extending along the coast of Oman, to compare the float trajectories with the known circulation. Floats were selected for our region of interest, in the Gulf of Aden up to the Ra's Al Hadd cape, along the western coast of Oman, with time of measurements from 6 months to 2 years. In the following paper, five of them are presented, two in the Gulf of Aden, 2 in the western Arabian Sea and one o↵ coast. These floats were particularly relevant, in the presence of Red Sea Outflow Water and Persian Gulf Water and with strong variabilities in their trajectories. In order to understand these trajectories, we compared them with the sea level anomaly recorded by AVISO, and with the bathymetry; we also focused on their distribution in kinetic energy. The results are presented hereafter in an article published on March 2012 in Ocean Science.

Introduction

The Northwestern Indian Ocean is composed of the 3000-5000 m deep Arabian Basin bounded east by the shallower Chagos-Laccadive ridge and the Indian continental shelf and west by the shallower Gulf of Aden. The Arabian Sea is bounded south by the Carlsberg Ridge (see Fig. 1). Southwest of the Arabian Sea, the Somali Basin has the largest depths (deeper than 5000 m; Tomczak and Godfrey, 2003). This oceanic region is connected to two evaporation basins, the Persian Gulf (with depths shallower than 90 m and mean depth of 25 m), and the Red Sea (with maximum depths near 2740 m and mean depth of 490 m). These two basins produce salty waters which flow into the Arabian Sea (see below).

The circulation and water masses in the Arabian Sea are strongly influenced by the monsoon winds and by the intense air-sea heat fluxes (Johns et al., 2000). The winter monsoon winds blow to the Southwest from December to April. In the Arabian Sea, the surface oceanic currents form a cyclonic gyre in January, which weakens in March [START_REF] Kindle | A review of the surface circulation of the Northern Arabian Sea[END_REF]. During the winter monsoon, the mixed layer deepens strongly offshore, a priori due to convective overturning in mid-basin, while the dominant Ekman transport is shoreward [START_REF] Lee | The upper-ocean respecton to monsoonal forcing in the Arabian Sea: seasonal and spatial variability[END_REF]. The summer monsoon winds blow to the Northeast from June to October. They are usually stronger than the winter monsoon winds and lead to faster currents. During the summer monsoon, a swift northeastward current (the East Arabian Current) flows south of the Arabian Peninsula, and forms the northern part of the regional anticyclonic gyre.

During the summer monsoon, upwelling occurs on the Southern coast of Oman, near Ras al Madrakah-Ras ash Sharbatat (Shi et al., 2000). Moreover, between early May and end of October, a jet-front system forms at Ras al Hadd Published by Copernicus Publications on behalf of the European Geosciences Union.

Abstract

(southeast of Oman), as cool upwelled waters of the Southern Omani coast meet the warmer waters of the Gulf of Oman [START_REF] Böhm | The Ras al Hadd jet: Remotely sensed and acoustic Doppler current profiler observations in 1994-1995[END_REF]. The associated southeastward jet is confined between two counter-rotating eddies (one cyclonic to the north and an anticyclonic one to the south).

During the same period, the northeastward Somali Current is also strongly intensified [START_REF] Schott | Winter Monsoon circulation of the Northern Arabian Sea and Somali Current[END_REF]. It is then accompanied offshore by a large anticyclonic eddy, the "Great Whirl". During summer, the monsoon winds also generate upwelling along the Somali coast. The cold and fresh upwelled waters (with temperatures of 17-18 C) are in strong contrast with the oceanic surface waters (temperatures of 28-29 C). Note that upwellings are also observed along the southwestern coast of India during the summer monsoon.

Furthermore, the Arabian Sea is a region of strong evaporation (the E P budget is on the order of 150 cm yr 1 ). Thus, surface salinity as high as 36 can be found in its northern part.

The circulation in the Gulf of Aden, and in the Gulf of Oman, is more complex, formed of local cyclonic and anticyclonic gyres (Bower et al., 2002;Fratantoni et al., 2006;Bower et al., 2000;Pous et al., 2004b). In the Gulf of Oman, one or two large cyclonic gyres have been identified by these authors. In November, due to winds blowing along the axis of the Persian Gulf, and which also affect the Gulf of Oman, upwelling can appear along the Iranian coast in the Gulf of Oman (and in the Persian Gulf, Pous et al., 2004a, b). In the Gulf of Aden, the Somali Current Ring and the Gulf of Aden Eddy appear at the end of the summer and winter monsoons respectively, and the Gulf of Aden eddy propagates westward into the Gulf while the Somali Current Ring remains relatively stationary at the entrance of the Gulf. Recently, [START_REF] Bower | Mesoscale Eddies in he Gulf of Aden and their Impact on the Spreading of Red Sea Outflow Water[END_REF] have shown the existence of two other eddies in the Gulf of Aden: the Lee Eddy and the Summer Eddy which form early during the summer monsoon due to strong winds blowing through the Socotra Passage.

The Northwestern Indian Ocean is noticeably influenced by the Persian Gulf and by the Red Sea. These two adjacent seas produce warm and salty waters (the Persian Gulf Water, PGW and the Red Sea Outflow Water, RSOW) which penetrate into the Northwestern Indian Ocean via the Gulf of Oman, and the Gulf of Aden, where they equilibrate at depths of 200-300 m and 600-1000 m respectively (Bower et al., 2000;Pous et al., 2004b;Bower et al., 2005). The Red Sea Water enters the Gulf of Aden with a temperature close to 22 C and a salinity close to 40. These values decrease to 16 C and 36 once this water has mixed with surrounding waters (Indian ocean Central Water, ICW) and the RSOW settles at depths of 600 to 900 m in the Arabian Sea; from there, it follows the East African coast southward. Near 30 S, the RSOW is found at depths of 1000 to 1100 m. The Persian Gulf Water has similar temperature and salinity values (22 C and 40) in the northern part of the Straits of Hormuz. Strong mixing with Indian Ocean Surface Water (IOSW) decrease the thermohaline characteristics of the outflowing water to about 20 C and to 37.5 in the Gulf of Oman where the equilibrium depth of PGW is about 200-250 m; the PGW characteristics decrease to 17 C and 36.2 in the northern Arabian Sea where its equilibrium depth is about 300 m (Prasad et al., 2001).

The other water masses in the region are the ICW, which originates from the Subtropical Convergence and lies in and above the thermocline (Tomczak and Godfrey, 2003). Its temperature range is 17 C to 5 C from top to bottom, and its salinity varies from 35.5 to 34.8. It also bounds the warm and salty outflow water masses of the Northern Arabian Sea on their Southeastern side.

Below the ICW, lies the Indian Deep Water (IDW, at depths between 1500/2000 m and 3800 m, with potential temperature of 2.5 C and salinity of 34.8) and finally the Antarctic Bottom Water (AABW), below 3800 m depth, with potential temperature of 0.3 C and salinity of 34.7.

In this paper, we will investigate the regional and mesoscale variability of the upper water masses (above 2000 m depth) with hydrological data gathered by ARGO floats. ARGO is an international program for the monitoring of the global ocean, which relies on hydrological data collection by about 3000 profiling floats.

These floats are released at the surface at different locations, dive to their 1000 m "parking depth" and most often remain there for 10 days (a few floats for 5 days only). Then they dive to 2000 m depth and perform a vertical profile of temperature and salinity up to the surface. When the floats surface, they transmit their data to the ARGO centers via satellites. This cycle is then repeated for the lifetime of the float. Once processed, the hydrological data can be mapped along the float trajectories, providing Lagrangian sections of temperature and salinity.

The paper is organised as follows: the dataset is presented and the floats retained for analysis are classified regionally. The long and short-term displacements of the floats are related to the variations of water masses and to the possible presence of mesoscale anomalies. Deep mesoscale displacements are studied in relation with surface motion provided by altimetric maps.

Data and methods

ARGO float data were obtained from the ARGO Regional Center for the Indian Ocean (INCOIS, Indian Government). In the database, 43 floats were identified with a mission beginning north of 12 N (i.e. a surface release north of this latitude). Among them, eight floats were discarded because of too short missions. Two floats were discarded because of erroneous hydrological data. Two floats evidenced positioning or date problems (showing jumps). Finally, thirteen floats were not analysed since most or all of their trajectory lay east of 65 E, while this paper concentrates on the Arabian Sea between the Gulf of Oman and the Gulf of Aden 20 float trajectories and hydrological datasets were studied.

Here, the analysis of 5 floats is presented. When a float performed a multi-year mission, its data were segmented into yearly bins starting at the date of float release. Temperature and salinity were plotted along the trajectories, as vertical sections.

To evidence correlations between the float trajectories and local bathymetric anomalies, ETOPO1 bathymetric maps were extracted from the NGDC database.

To establish relations between motions at 1000 m depth and surface motions, altimetric maps were obtained from the CNES/AVISO satellite data center. Delayed-time sea-level anomalies (SLA, with 1/3 ⇥1/3 resolution on a Mercator grid and with daily interpolation), computed with respect to a 7-year mean of the sea surface topography, are used here to determine the correlation with the float motion.

For each float, we shall show the vertical/alongpath section of temperature and salinity, and we shall try to relate the major thermohaline anomalies with the water masses described in the introduction. We shall also plot the trajectories superimposed on the bathymetry or on SLA. To establish quantitative correlations between the float trajectories and the bathymetry, we shall compute the angle between the trajectory and the isobaths at each position, and plot this angle versus time.

For each float, we shall compute Lagrangian velocities from the distance and time-lag between two surfacings; these surfacings can be consecutive, or separated by two cycles or more. We shall plot the kinetic energy distribution thus obtained with respect to the distance covered by the float, or with respect to the orientation of its trajectory. Note that the orientation is given between -90 and +90 , with the origin on the eastward axis. The direction of the flow (e.g. eastward or westward for a zero degree angle) is here less important. For instance, energy in both northeastward and southwestward directions reflects the influence of monsoon currents.

Also, our analysis of the floats energy assumes that the flow field has global homogeneity, isotropy and stationarity. These hypotheses will be tested by computing the energy distributions for several floats (thus having different starting points and different trajectories), and for different periods of the year.

Gulf of Aden

Float 1900432

Float 1900432 was released 8 April 2007 near 45 52 0 E, 11 58 0 N, and its mission ended on 17 October 2007 near 46 40 0 E, 13 24 0 N. We terminated the analysis of its data on the 23 September 2007 since the last hydrological measurements were not realistic. The 4-day cycle of the float allowed the observation of short-term loops.

Its motion, shown on figure 2 was mostly cyclonic. Firstly, the trajectory was northeastward from about 46 N, 12 E to 46 30 0 E, 12 30 0 N for 20 days. Then the float performed one large cyclonic loop with diameter of about 100 km, for 40 days, in the area spanning 46 30 0 E to 47 30 0 E and 11 50 0 N to 13 05 0 N; this large loop was followed by two small cyclonic loops at 12 30 0 N near 46 40 0 E and 45 50 0 E, between days 56 and 100 of the mission. Then a larger loop (with diameter of about 100 km) was performed cyclonically until day 128 (in the area between 12 N and 13 N and between 46 E and 45 E). Finally, the float drifted northeastward to 48 E, 13 N and then northwestward to 47 30 0 E, 13 30 0 N. Velocities ranged between 4 and 20 cm s 1 . The kinetic energy distribution with respect to horizontal displacement is shown in Fig. 3; this distribution shows four relative maxima (we consider that a wide peak in the curve, or two peaks very close to each other constitute a single maximum). The largest kinetic energy (K ⇡ 0.03 m 2 s 2 ) is found at horizontal scales of 40 km, most likely, in relation with the two small loops with swift motion of the float; a second relative maximum is formed of two peaks, at 60 and 80 km horizontal scales; this can be related to the fast and small zonal motions of the float and to the cycloidal motion superimposed on the large loops in its trajectory; the third relative maximum occurs for 120 km horizontal scales, in relation with the large loops themselves. Finally, the peak at 170-190 km can be associated with the large zonal motion of the float. In the large loops, the float velocity was about 5 cm s 1 . This value is in agreement with Bower and Furey's findings (2011): the velocities in the RSOW are weaker in the western Gulf of Aden (5 cm s 1 ) than in the eastern gulf (15 cm s 1 ). Unfortunately, the time-life of the float was too short to test the invariance of this spectrum to the period of the year (this point will be addressed in more details in Sect. 5 of the paper). One can only state that the larger scale motions occured at the end of the mission, that is, in late summer, early fall.

This figure also shows the energy distribution versus the angle of the flow. This distribution has two peaks at = +25 and 15 corresponding to the along-gulf displacements, and at = +65 and 75 corresponding to acrossgulf displacements. Obviously, energy is also found at other angles (e.g. at = 35 ) since several loops were present in the trajectory.

The float motion superimposed on the bathymetry is shown in Fig. 2. The angle between the trajectory and the isobaths was computed and plotted versus time (figure not shown). This angle was often larger than 45 ; it was smaller than 20 only at days 80-90 and 110-120. This confirms that the correlation between isobaths and float trajectories was weak except close to the continental slope.

The float motion is now compared with the surface flow, given by the SLA. anticyclonic west of this longitude, were present in the SLA map (see Fig. 4). The float trajectory, superimposed on this SLA map, indicates that part of the large loop performed by the float from days 20 to 60, circled the cyclonic gyre. The small indentation in the trajectory between days 32 and 40 of the mission (8 May and 16 May) may be due to a positioning error, or to small scale flow not sampled by altimetry.

On 27 June (day 80 of the mission), the surface circulation was cyclonic west of 47 E and anticyclonic east of that longitude. The float trajectory nearly circled the core of the cyclonic gyre. One month later (25 July or day 108 of the mission), the anticyclonic gyre had retreated to the southeast and the flow was cyclonic over most of the region; again the float circled anticlockwise around the cyclonic region and then followed the southwestward flow between the cyclone and the anticyclone.

Thus, the float motion was consistent with the surface circulation indicated by the SLA. Indeed, gyres in the Gulf of Aden have a noticeable barotropic component (Bower et al., 2002;[START_REF] Bower | Mesoscale Eddies in he Gulf of Aden and their Impact on the Spreading of Red Sea Outflow Water[END_REF].

The water masses sampled by the float during its vertical profiles are now identified on the vertical sections of temperature and of salinity along the trajectory of float 1900432 (see Fig. 5).

From day 0 to 24 of the mission, the float encountered a warm and salty water mass, at depths between 400 m and 1000 m, with temperature ranging between 12 C and 16 C and salinity between 35.8 and 36.5. The maximum temperature was slightly shallower than the maximum salinity (at 400-600 m versus 500-700 m). This water mass was clearly Red Sea Outflow Water. Very warm and salty waters were again found with peaks between days 80 and 100, and then at days 130-140. Thus RSOW was mostly confined west of 47 E. Deep data were missing at days 108-116 and no hydrological data were returned at days 120 to 128, hence the gap observed in the temperature and salinity sections.

On the contrary, east of 47 E, cooler and fresher water was found in the 300-800 m depth range, with temperature lying in the 9-13 C interval and salinity between 35. 5 and 35.8. This water was more likely ICW, mixed with RSOW, to account for the slightly larger salinity than that of pure ICW. This water was contained in the large cyclone which lay around 47 E, 12 30 0 N at day 40 of the mission. This confirms the conclusion drawn by [START_REF] Bower | Mesoscale Eddies in he Gulf of Aden and their Impact on the Spreading of Red Sea Outflow Water[END_REF] that westward drifting eddies can transport ICW into the Gulf of Aden and mix it with RSOW there.

Finally, the thermocline waters became thinner during the 6 months of the record, either because of wind-induced vertical mixing, or of lateral turbulent mixing, or because of the presence of cooler water close to the northern coast. of the region; again the float circled anticlockwise around the cyclonic region and then followed the southwestward flow between the cyclone and the anticyclone. Thus, the float motion was consistent with the surface circulation indicated by the SLA. Indeed, gyres in the Gulf of Aden have a noticeable barotropic component (Bower et al., 2002;[START_REF] Bower | Mesoscale Eddies in he Gulf of Aden and their Impact on the Spreading of Red Sea Outflow Water[END_REF].

Float 2900392

The water masses sampled by the float during its vertical profiles are now identified on the vertical sections of temperature and of salinity along the trajectory of float 1900432 (see figure 5).

From day 0 to 24 of the mission, the float encountered a warm and salty water mass, at depths between 400 m and 1000 m, with temperature ranging between 12 o C and 16 o C and salinity between 35.8 and 36.5. The maximum temperature was slightly shallower than the maximum salinity (at 400-600 m versus 500-700m). This water mass was clearly Red Sea Outflow Water. Very warm and salty waters were again found with peaks between days 80 and 100, and then at days 130-140. Thus RSOW was mostly confined west of 47 o E. Deep data were missing at days 108-116 and no hydrological data were returned at days 120 to 128, hence the gap observed in the temperature and salinity sections. On the contrary, east of 47 o E, cooler and fresher water was found in the 300-800 m depth range, with temperature lying in the 9 -13 o C interval and salinity between 14 initiated, this float was still active. Only the first two years of its mission are presented here (in two segments of one year each). This float had a 5-day cycle.

During year 1, the float drifted eastward for a month from 46 13 0 E to 48 15 0 E along 11 40 0 N (see Fig. 6). Then, it followed three loops, one anticyclonic, one cyclonic and then a large anticyclonic one at 49-50 E, and later it moved up to 51 E, 14 40 0 N. Finally it came back swiftly southwestward along the coast of Yemen (again with a cyclonic loop). Clearly, the float trajectory was a combination of motion around gyres and of fast along-gulf displacement near the northern shore (during the fall intermonsoon).

Only a weak correlation of the float trajectory with the isobaths was found, except in the last stage of the float evolution (between days 330 and 350 when the angle between the trajectory and the isobaths was smaller than 10 ).

In the energy distribution (see Fig. 7), a wide peak at 30-40 km horizontal scale corresponds to the small scale loops in the float trajectory, two peaks at 60 and 90 km can be related to the meanders in the trajectory, and we associate smaller peaks at 200 and 260 km with the large excursions of the float along the gulf axis. In the angular distribution of the energy, two peaks are found at -15 and +25 (as for float 1900432); this corresponds to displacements mostly along the Gulf axis. A peak at a -45 angle can be related to the swift float motion at days 120-135, 160-200 and 300-340. Altimetric data were used to determine the surface circulation (see Fig. 8). On day 115 of the mission (25 January 2006), a strong cyclonic gyre was present near 48 30 0 E and 12 30 0 N; this gyre had a tail on its southeastern side. The float trajectory was anticlockwise and followed this tail. On day 160 of the mission (8 March 2006), the circulation near 49 30 0 E and 13 N was strongly anticyclonic; the float trajectory was clockwise and then followed the southeastward flow between this anticyclonic gyre and the cyclonic one on its eastern flank. Finally, on day 325 of the mission (23 August 2006), the regional gyre circulation was again cyclonic over the central part of the Gulf of Aden. The float trajectory was clearly anticlockwise around this gyre. This confirms the deep influence of the regional gyres observed in the altimetry.

The thermohaline section along the float trajectory is displayed in Fig. 9. During the first 100 days, the float sampled warm and salty RSOW. Again, the temperature and salinity maxima were about 16 C and 36.5, located respectively at depths between 400 and 500 m and between 500 and 600 m. This thermohaline anomaly diminished as the float moved eastward, in particular after day 120 when the trajectory exhibited the large anticyclonic loop. Clearly, RSOW lay in the southwestern part of the Gulf of Aden at that time. Then, the temperature and salinity measured by the float at 600 m depth were on the order of 12 C and 35.8. Again, these values were slightly above ICW characteristics, indicating mixing with RSOW. The surface layer became cooler and fresher in summer and warmer and saltier in winter when the winds were weaker (an observation similar to that of float 1900432).

Ocean

During year 2, the float followed a series of cyclonic and anticyclonic meanders, with a large cyclonic loop around 48 E, 12 30 0 N for about 30 days (see Fig. 10). Then, the float moved rapidly northeastward for 20 days and performed a cyclonic loop between 51 E and 52 E, and 12 30 0 N and 13 30 0 N, for about 40 days. Then the float drifted back southwestward, nearly to its starting point in less than 4 months. From there, it followed three meanders, alterna- tively cyclonic, anticyclonic and cyclonic, up to 50 E, 14 N and finally drifted back southwestward again. This trajectory illustrates the combination of fast along-gulf currents, especially near its northern coast (with a reversal of direction between winter monsoon and spring intermonsoon), and of large cyclonic and anticyclonic gyres at the center of this gulf.

Correlation of the float motion with the isobaths orientation was again weak, except close to days 450 and 550 when the float moved along the northern coast.

The kinetic energy distribution (not shown) has peaks at 40, 80, 140 and 400 km, corresponding respectively to the small and medium size gyres or meanders, for the first three, and to the fast along gulf motion at days 500-600, for the latter. The energy distribution versus the angle has maxima at +20 and -80 , which are the along-gulf and across-gulf orientations.

The altimetric maps agree with this circulation (see Fig. 11): in November 2006, the surface gyre was cyclonic above the first cyclonic loop of the float. In January 2007, the float first moved clockwise around the anticyclone at 50 E, 13 30 0 N; then the float performed a cyclonic loop at 1000 m depth under a cyclonic surface gyre, located near 52 E, 14 N. Finally, a clear correlation can also be established between the anticyclonic loop of the float near 48-50 E and 13-13 30 0 N, and an anticyclonic surface gyre in July 2007.

Again (see Fig. 12), thermohaline sections along the float trajectory show that warm and salty water lay in the western part of the Gulf around cyclonic gyres a depths close to 600 m. This water had the same thermohaline maxima as in year 1 (about 16.5 C and 36.5 at respectively 500-600 m depth and 600-700 m depth), and thus was a priori RSOW. Colder and freshed water (with temperature of 10-12 C and salinity of 35. 5-35.8) was sampled by the float at days 500-550 and 650-700, in the northeastern part of the gulf. Again this water was ICW mixed with RSOW by the large gyres. During the first year, the float drifted northeastward to 56 E, 13 N and then northwestward to 54 30 0 E, 14 N within about 100 days (see Fig. 13). It remained stationary near this position for about two months, then it performed a cyclonic loop between this position and 56 E and 15 N. After day 250, it followed a large anticyclonic loop (with strong zonal predominance) up to 58 E and to 17 30 At depths between 300 and 800 m, a core of cold and fresh water was sampled by the float between days 0 to 20, 180 to 210 and 260 to 270. With temperatures of 8 to 12 C and salinities of 35.1 to 35.4, this was likely ICW.

At depths between 400 and 900 m, two cores of warm and salty Red Sea Outflow Water were present from day 30 to 60, and from day 140 to 160. The first warm core was located east of Socotra, near the continental slope; this core of RSOW flowed around the island. The salinity and temperature at 700 m depth were about 35.9 and 12 C. Similar maxima were found again near day 160, at the same depth, as the float was much farther north (near 55 30 0 E, 14 10 0 N). At that time, the float performed a cyclonic loop, and a cold core was sampled on the northern part of this loop. We interpret these two events as the advection of a detached fragment of RSOW by a cyclonic gyre.

Finally, in January-February 2007, a weak signature of Persian Gulf Water (T ⇠ 18 C, S ⇠ 35.9 at depths near 200-300 m) was sampled by the float, as it came near Ras al Madrakah (near 57 50 0 E, 19 N).

In the altimetric map, a cyclonic surface signature was present between days 180 and 210 when the float followed a cyclonic trajectory near 55 30 0 E, 14 20 0 N (see Fig. 15). Early January 2007, the anticyclonic loop of the float near 57 E, 16 N was also related to an anticyclonic sea level anomaly. Though the correspondence between float trajectory and SLA was not as strong as in the Gulf of Aden, the deep motion of the float could still be correlated with the surface motion at horizontal scales of 100 to 200 km, in this region.

During year 2, float 2900104 performed successively a cyclonic, an anticyclonic and a cyclonic loop, for about 40 days each, in a general northeastward direction (towards Ras ash Sharbatat (near 57 E, 18 N; see Fig. 16)). After about 150 days, the float followed a cyclonic loop northwestward towards this cape (for about 3 months), and then an anticyclonic loop northeastward for about two months. The end of its trajectory (for year 2) was southwestward. The energy peaks (not shown) are located at 50, 80, 120, 160 km horizontal scales, and with a smaller amplitude, at 220, 270, 320 and 410 km. The former four peaks correspond to the float motion between days 500 and 650 while the latter four correspond to the larger scale displacements near the beginning and end of this second year of mission (days 365-500 and 650-730). The energy distribution with respect to the orientation (not shown) peaks by far at a 25 angle, corresponding to the large scale ENE-WSW displacements of the float with the monsoon currents.

Only a weak correlation was found between this float trajectory and bathymetry, except at days 400-450 and 600-650, but mostly over the abyssal plain. In fact, the correlation over the abyssal plain may be due to regional gyres (at 100 km scale), which remain east and west of Ras al Madrakah for long periods of time. When a float is trapped in these gyres, it remains over the abyssal plain and does not cross isobaths.

For this float, surface-depth correlations of motions were found at horizontal scales of 200 km, but those at 50 km scale were not always clear. At larger horizontal scales, (400-500 km), the float trajectory reflected the monsoon currents.

The temperature and salinity data collected by this float during year 2 show evidence mostly of Persian Gulf outflow Water, especially near days 500, 600 and 680 (about end of July, early November 2007 and mid January 2008; see Fig. 17). This water mass was thus found east of Ras ash Sharbatat and south of Ras al Madrakah, for the first two periods, and south of Ras ash Sharbatat for the latter. The stronger salinities for PGW were found in November and in January.

This indicates that (a) PGW can flow southwestward downstream of Ras ash Sharbatat, though most likely as fragments, (b) this flow intensifies during the transition from summer to winter monsoon, and during winter monsoon, as the large-scale, wind-induced circulation is southwestward. Similar maxima were found again near day 160, at the same depth, as the float was much farther north (near 55 o 30 E, 14 o 10 N). At that time, the float performed a cyclonic loop, and a cold core was sampled on the northern part of this loop. We interpret these two events as the advection of a detached fragment of RSOW by a cyclonic gyre. Finally, in January-February 2007, a weak signature of Persian Gulf Water (T ∼ 18 o C, S ∼ 35.9 at depths near 200-300 m) was sampled by the float, as it came near Ras al Madrakah (near 57 o 50 E, 19 o N).

In the altimetric map, a cyclonic surface signature was present between days 180 and 210 when the float followed a cyclonic trajectory near 55 o 30 E, 14 o 20 N (see figure 15). Early January 2007, the anticyclonic loop of the float near 57 o E, 16 o N was also related to an anticyclonic sea level anomaly. Though the correspondence between float trajectory and SLA was not as strong as in the Gulf of Aden, the deep motion of the float could still be correlated with the surface motion at horizontal scales of 100 to 200 km, in this region.

During year 2, float 2900104 performed successively a cyclonic, an anticyclonic and a cyclonic loop, for about 40 days each, in a general northeastward direction (towards 27 Colder and fresher waters were also identified at 200 depth around day 550 (early September 2007) and day 620 (mid November 2007). Indeed, in September, upwellings can develop near Ras ash Sharbatat and bring colder and fresher waters upward.

On the contrary, in November, isotherms rose above 300 m depth, but dived below 600 m depth. This structure of the isotherms, the values of temperature and salinity at depths between 300 and 600 m, and the anticyclonic loop in the float trajectory, point towards the existence of a deep anticyclone of ICW at that time and location. This anticyclone may have been generated earlier by the instability of the northeastward current along the southern coast of Oman (see also [START_REF] Lee | The upper-ocean respecton to monsoonal forcing in the Arabian Sea: seasonal and spatial variability[END_REF].

Float 1900438

Float 1900438 was launched 18 November 2008 at 56 20 0 E, 16 04 0 N with a 4-day cycle, and its validated data ended 14 March 2009, at 54 26 0 E, 16 02 0 N. After that date, a 5 month long data gap occurs corresponding to a shift in position by 7 degrees east.

Its initial trajectory was mostly southwestward, in accordance with the monsoon current direction (see Fig. 18). On this global drift, two cyclonic loops were superimposed, one between days 35 and 50 and near 55 E, 14 N, the other one between days 80 and 100 and near 52 30 0 E, 14 N. After 100 days, the float drifted back to the Northeast along the coast of Oman, perhaps in relation with an early reversal of the monsoon current.

The energy distribution with distances (not shown) has peaks at 30, 100 and 150 km horizontal scales, in agreement with the small and large loops and with the ENE-WSW displacements. In the angular distribution, the energy maxima lie at +30 , -35 , +65 and -75 . The first orientation is clearly related to the large-scale motion with the monsoon currents.

The angle between the trajectory and the isobaths (see Fig. 19) was most often small because the float crossed few topographic anomalies or flows along the continental slope. Nevertheless, the bottom topography did not seem to steer the float motion, but the currents with surface expression, or the anomalous water masses, did (see below).

The temperature and salinity sections of this float show the presence of PGW at 200 m depth during the first 20 days of the mission, south of Ras ash Sharbatat (near 57 E, 18 N); this was most likely a fragment detached from the PGW outflow (see Fig. 20). Between days 32 and 50 (during the first loop in the trajectory), no deep thermohaline maximum appeared. But a noticeable RSOW signature was visible during the second loop and once the float drifted eastward, along the continental slope, towards 54 40 0 E, 16 N.

Again, there was a good correlation of the float trajectory with the SLA (see Fig. 21). Both cyclonic loops circled around surface depressions (days 36 and 72 of the mission) and the final clockwise motion of the float wrapped around a surface elevation (days 92 of the mission).

Float 2900394

This float was released on 27 June 2005 at 62 33 0 E, 20 50 0 N but a positioning problem occured for the first 10 days. Then, the trajectory was correct, but a major drift occured in temperature data for 40 days. Therefore the data were analysed only from day 50 onward. This float had a 5-day cycle.

Float 2900394 performed a series of anticyclonic loops between 63 30 0 E and 64 30 0 E and 19 30 0 N and 21 N for about 240 days starting from day 50 (see Fig. 22). The abrupt shift to the south between days 175 and 185 may be a positioning problem. Then, to the end of year 1, the float followed an anticyclonic trajectory from 64 E, 20 N to 65 30 0 E, 19 N and back to 63 30 0 E, 18 30 0 N.

The angle between the trajectory and the bathymetry varied rapidly with time, and did not indicate a durable correlation. The energy distribution with distance indicates a peak at 25 km, related to the small loops performed by the float (not shown). Its distribution with respect to the angle has peaks at ±85 , related to the meridional motion. Other peaks at ±20 and ±40 have smaller amplitudes.

The thermohaline anomalies identified during this period were located between 150 and 350 m depths (see Fig. 23). Warm and salty water (17-18 C, 36.2-36.4) was observed at days 50-100, 150-240 and 270-320. This was clearly PGW and it was concentrated near 20-21 N, 63 30 0 E-64 30 0 E. Considering the float motion, this may have been an eddy of PGW. On the contrary, a cold and fresh anomaly was sampled around day 340. This water may have originated from the Ras al Madrakah upwelling which was active then.

The altimetric maps during the September 2005-April 2006 period, in the 63-65 E, 18-21 N region, did not show durable correlations with the float trajectory, except during the last two months. This suggests that the anticyclonic motions of the float during its first eight months were more related to the deep, warm and salty water masses, that it encountered.

During the second year, the trajectory of float 2900394 was mostly southeastward, with one small anticyclonic loop near 64 30 0 E, 18 N, and two cyclonic loops near 64 30 0 E, 17 N and 65 E, 17 30 0 N (see Fig. 24). This repeated southeastward drift of float 2900394 indicates the presence of a current in this direction, in summer, from the mouth of the Gulf of Oman. This current is a priori the Ras al Hadd jet. The correlation with altimetry was more durable for the large cyclonic motions. The vertical sections of temperature and of salinity showed again three blobs of PGW at days 420-510, 550-600, 650-680 (see Fig. can develop near Ras ash Sharbatat and bring colder and fresher waters upward. On the contrary, in November, isotherms rose above 300 m depth, but dived below 600 m depth. This structure of the isotherms, the values of temperature and salinity at depths between 300 and 600 m, and the anticyclonic loop in the float trajectory, point towards the existence of a deep anticyclone of ICW at that time and location. This anticyclone may have been generated earlier by the instability of the northeastward current along the southern coast of Oman (see also [START_REF] Lee | The upper-ocean respecton to monsoonal forcing in the Arabian Sea: seasonal and spatial variability[END_REF]. Its initial trajectory was mostly southwestward, in accordance with the monsoon current direction (see figure 18). On this global drift, two cyclonic loops were superimposed, one between days 35 and 50 and near 55 o E, 14 o N, the other one between days 80 and 100 and near 52 o 30 E, 14 o N. After 100 days, the float drifted back to the Northeast along the coast of Oman, perhaps in relation with an early reversal of the monsoon current.

Float 1900438

The energy distribution with distances (not shown) has peaks at 30, 100 and 150 km 31 interval, a small anticyclonic loop was achieved by the float; two cyclonic loops were performed during the two other periods.

Discussion

Here, we discuss two points which concern all floats: the dependence of the spatial distribution of kinetic energy with respect to location or time, and the correlation of temperature and salinity measurements by the floats with the regional climatology. Firstly, we study the homogeneity, isotropy and stationarity of the kinetic energy distributions.

In the Gulf of Aden, these distributions for the two floats concerned, for each year, and even for each season, showed the following: The distribution with respect to the distance had three peaks, at 30, 60 and 160 km, this latter being slightly less marked during the summer. The annual distributions with respect to the angle had peaks at ±20 , +60 , -40 , and -80 , while the seasonal distribution had peaks at +20 , +60 in summer and at -20 and -80 in winter.

From this we can conclude that the same scales of motion were present in both seasons and at different locations in the Gulf, mostly due to the gyres which occupied it. Though our database is limited, so that conclusions are difficult to draw for the orientation of the trajectories with respect to seasons, we observed twice a westward-southwestward current (angle of +20 ) along the coast of Yemen at the end of the winter monsoon. This can be related with the presence of cyclones at this location, already described by [START_REF] Bower | Mesoscale Eddies in he Gulf of Aden and their Impact on the Spreading of Red Sea Outflow Water[END_REF].

In the Arabian Sea, the peaks in the distribution of kinetic energy with respect to distance were found at 30, 60, 100 and 150 km, except for float 2900394 for which the energy was concentrated by far at the 20-25 km scale. The peaks at scales of 60 and 100 km were related to gyre motion.

Concerning the orientation of motion, the energy peaks at +20 to +40 were present for all floats. This peak was usually related to large scale displacements on seasonal time scales (related to monsoon currents). Other peaks were related to gyre motions or to local currents such as the Ras al Hadd jet.

Therefore, there was similarity in the dominant scales of motion between floats in the Arabian Sea, but also between the Arabian Sea and the Gulf of Aden.

Secondly, we compare the vertical sections of temperature and salinity obtained by the floats to the climatological temperature and salinity maps at 300 and 600 m depth (the depth of PGW and RSOW or ICW). These maps are given in appendix (see Figs. A1 andA2).

In the Gulf of Aden, the climatological maps show a zonal gradient of temperature and salinity, which intensifies in May. This gradient is shifted eastward and is weaker in November. During the 160 days of its mission, float 1900432 measured a maximum thermohaline gradient near 46 30 0 E as in the climatology. Float 2900392 evolved mostly east of this longitude and was thus located in a region of weaker salinity anomaly. From January to May, the climatology indicates that the warm and salty water at 600 m depth and east of 46 30 0 E, extends eastward along the coast of Yemen. Later in the year, this eastward extension occurs rather along the coast of Somali. For years 1 and 2, float 2900392 resided more often in the northern part of the Gulf of Aden. Thus it showed stronger salinity anomalies from September to March than later in the year. During year 2, float 2900392 showed salinity maxima at days 400 and 600 at its westernmost position in the Gulf (near 47 E).

In the Arabian Sea, a positive salinity anomaly is observed in the climatology north of Socotra, at 600 m depth; this anomaly is most marked in July. Float 2900104 also indicated a salinity anomaly near 14 N, 54-56 E at this depth and in May-September (year 1). The same float measured a PGW anomaly near Ras al Madrakah in mid January (year 1); the climatological salinity map at 300 m also shows a wide extent of the PGW plume at this location in January (see Fig. A). Float 2900104 showed again strong salinity anomalies at 300 m depth in early August, mid November and in January (year 2). Climatology indicates that this is the period of maximum extent of PGW along the southern coast of Oman. PGW was also measured by float 1900438 off Ras ash Sharbatat in November; this corresponds to the large extent of PGW in the climatology at that time. The same float sampled RSOW near 53 E and 14-15 N in January, again at a time of wide extent of RSOW along the coast of Yemen, in the climatology. Finally, float 2900394 sampled PGW near 64 E, both during year 1 (near 20-21 N) and during year 2 (near 16-17 N), except in March-May. Climatology indicates that PGW extends south of 20 N at 64 E except in March-May.

Conclusions

The data from five profiling floats in the Gulf of Aden and in the Arabian Sea have been analysed. Their trajectories and thermohaline anomalies at depth have been compared with surface motions, with bathymetry and with climatological water masses. Clearly, if a strong correlation between surface and deeper motions could often be established at 100-200 km scale, loops with 30-40 km diameter, performed by the floats were not sampled by altimetry, due to its resolution. Some of these small loops could be correlated with thermohaline anomalies at the same scale and at the depth of PGW or of RSOW. The float trajectories also reflected the large-scale wind-induced circulation. Firstly, we study the homogeneity, isotropy and stationarity of the kinetic energy distributions.

Ocean

www

In the Gulf of Aden, these distributions for the two floats concerned, for each year, and even for each season, showed the following: The distribution with respect to the distance had three peaks, at 30, 60 and 160 km, this latter being slightly less marked during the summer. The annual distributions with respect to the angle had peaks at ±20 o , +60 o , -40 o , and -80 o , while the seasonal distribution had peaks at +20 o , +60 o in summer and at -20 o and -80 o in winter.

From this we can conclude that the same scales of motion were present in both seasons and at different locations in the Gulf, mostly due to the gyres which occupied it. Though our database is limited, so that conclusions are difficult to draw for the orientation of the trajectories with respect to seasons, we observed twice a westward-southwestward current (angle of +20 o ) along the coast of Yemen at the end of the winter monsoon.

This can be related with the presence of cyclones at this location, already described by [START_REF] Bower | Mesoscale Eddies in he Gulf of Aden and their Impact on the Spreading of Red Sea Outflow Water[END_REF].

In the Arabian Sea, the peaks in the distribution of kinetic energy with respect to distance were found at 30, 60, 100 and 150 km, except for float 2900394 for which the 39 In summary, the circulation in the Gulf of Aden is clearly composed of cyclonic and anticyclonic gyres at regional scale, observed in SLA, as mentioned in the literature, and with a barotropic character. The circulation in the Arabian Sea is also influenced by mesoscale gyres near the coast of Oman, which are well observed in altimetric data. More south, the circulation is complex and the float trajectories are less well correlated with SLA.

The scale and angular distribution of kinetic energy represents quite faithfully the different elements of the float trajectories. This distribution has common features from float to float or from year to year in a given region.

The correlation of the float trajectories with bathymetry was poor, except in one case; but this event may have been related to a quasi-stationary gyre located in a region of weak bathymetric anomalies.

Concerning the evolution of water masses, the surface temperature and salinity are strongly influenced by the seasonal variations of the wind, or by its sudden intensification. The only deeper-reaching influence of the wind observed by the floats was the appearance of colder and fresher water masses at 200 m depths, a priori related to upwelling events. Deeper cold and fresh water masses (e.g. between 500 and 800 m depths) were observed under the form of deep eddies or fragments of ICW. Concerning the orientation of motion, the energy peaks at +20 o to +40 o were present for all floats. This peak was usually related to large scale displacements on seasonal time scales (related to monsoon currents). Other peaks were related to gyre motions or to local currents such as the Ras al Hadd jet.

Ocean

Therefore, there was similarity in the dominant scales of motion between floats in the Arabian Sea, but also between the Arabian Sea and the Gulf of Aden.

Secondly, we compare the vertical sections of temperature and salinity obtained by the floats to the climatological temperature and salinity maps at 300 and 600 m depth (the depth of PGW and RSOW or ICW). These maps are given in appendix (see figures 26 and 27). In the Gulf of Aden, the climatological maps show a zonal gradient of temperature and salinity, which intensifies in May. This gradient is shifted eastward and is weaker in November. During the 160 days of its mission, float 1900432 measured a maximum thermohaline gradient near 46 o 30 E as in the climatology. Float 2900392 evolved mostly east of this longitude and was thus located in a region of weaker salinity anomaly. From January to May, the climatology indicates that the warm and salty water at 600 41 Intense thermohaline anomalies were measured at the depths of PGW or of RSOW.

In the Gulf of Aden, temperature and salinity peaks of 16 C and 36.5 were observed in the RSOW. The RSOW signal measured by floats intensified notably when the float circulated cyclonically in the western Gulf of Aden. This was well related to climatology also. The float recordings allowed us to identify the outflow of RSOW from the Gulf of Aden on the continental slope near 55 E and 13 N. It also showed evidence of detached fragments of RSOW north of that location (near 14 30 0 N).

In the Arabian Sea, temperature and salinity peaks of 19 C and 36.5 were observed in the PGW at 300 m depth. The PGW signal was observed near Ras al Hadd and Ras al Madrakah/Ras ash Sharbatat. This indicates that, under favorable conditions of monsoonal currents (in winter), the PGW outflow can veer around Ras al Hadd and extend southwest along the southern coast of Oman. Our observations also showed little PGW west of Ras ash Sharbatat. On the contrary, during summer, the PGW was exported southeastward from the Gulf of Oman; the Ras al Hadd jet was likely to focus this export in this direction.

Again, the seasonal variations of RSOW and PGW extent were observed in the float measurements in the Arabian Sea.

Further work should investigate this water mass distribution with more floats.

Appendix

Further discussion

Gulf of Aden

As underlined earlier the trajectories of the floats and the sea level anomalies show correlation and the high salinity water masses have been observed farther than on the climatological maps. The thermohaline sections also reveal variations of the measured water masses in function of the direction of the float. The two first floats in the Gulf of Aden (numbered 1900432 and 2900392) measured two di↵erent water masses, the highly saline Red Sea Outflow Water and the Indian Ocean Central Water during the same year, 2006 (second year of float 2900392). The Indian Ocean Central Water is observed on the large loops, on the eastern part of the Gulf of Aden, when the floats turns westward. On the other hand, the Red Sea Outflow Water is recorded by the floats with an eastward motion out of the basin. These section reveal di↵erent pathways of the water masses around or inside the gyres.

We stated in the beginning of this chapter that the use of few floats allowed us to have a qualitative study of the Gulf of Aden. We focus here on a quantitative analysis, using all the floats available for a specific period. On figures 3.1 and 3.2 we can compare the sea surface anomalies from the altimetry with the salinity and temperature measurements from over 250 measurements from ARGO floats in the Gulf of Aden, here at 600 m depth, where the peaks of Red Sea Outflow Water were observed previously.

In winter the maximum of temperature and salinity are observed along the northern coast of the Gulf of Aden with a recess near 47 o E, whereas fresher water are concentrated in the southern part. In the eastern part, diluted Red Sea Outflow water are found around 14.5 o N; 51 o E, with oceanic water penetrating the Gulf South of it. These positions correspond to the periphery of two strong anticylonic eddies observed on the altimetric anomaly map. Oppositely, in summer the highly saline and warm waters are observed along the southern part of the basin wheras the colder and fresher waters inflow along the coast of Yemen, with a front observed in the middle of the basin from 13.5 o N; 47 o E to 12 o N; 50 o E. The altimetry reveals a strong negative anomaly in the major part of the basin, with the exception of an anticylonic eddy west of the thermohaline front. This reinforces the conclusion that this highly seasonal circulation plays an important role of the pathways of the water masses via the observed eddies. This also underlines the fact that the GDEM climatology modified by ARGO reveals circulation with a mesoscale resolution.

In April 2006, during the spring inter monsoon, in the westernmost part of the Gulf of Aden, float 1900432 observed Red Sea Outflow Water with a westward motion along the coast of Yemen, whereas float 2900392 (second year) observed it in the southern part but with an eastward motion. During the spring inter monsoon, winds lose their intensity and main blowing direction, and the sea level anomaly signature is weaker (see figure 3.3). On a spring inter monsoon modified climatology, the saline waters are observed inside the eddy, contrary to the monsoon season where mixed water are found inside. This underlines that an intricate interaction between mesoscale gyres and water masses exists, but also that the dynamics are di↵erent during the inter monsoon. Hereafter, we focus on the Arabian Sea in order to observe if these interactions are concentrated in the Gulf of Aden or also prevail in the open ocean. Chapter3. Mesoscale variability of water masses in the Arabian Sea as revealed by Argo floats

54 o E 56 o E 58 o E 60 o E 62 o E 64 o E 66 o E o N o N o N o N o N o N Map of
E o N o N o N o N o N o N
Map of the 1757 Argo measurements distribution and their representativity ranges at 200 m depth during the summer monsoon 3.11. Further discussion

Arabian Sea

The floats on the Arabian Sea highlights also the seasonal variations of the gyres, via their trajectories. Float 2900104 leaves Socotra and remains blocked for several weeks at the mouth of the Gulf of Aden. Then it flows north-eastward with large loops along the coasts of the Arabian Peninsula, from the middle of the winter monsoon to the end of the summer monsoon. During these loops, the float also records patches of Persian Gulf Water. Finally if is quickly advected south-westward at the beginning of the winter monsoon, without any loops. The same winter monsoon advection goes for float 1900438 but with larger loops and then flows straight along the coast in early summer with a strong presence of Red Sea Water. The last float, 2900394, is farther from the coast, shows fewer loops and moves slower than the others, while recording Persian Gulf Water.

The surface circulation is influenced by the monsoon cycle, with sometimes, mesoscale eddies flowing south-westward or the Oman Costal Current with both north-eastward or south-eastward directions and a calmer o↵ shore circulation. On figures 3.4 and 3.5 the changes between monsoons are very clear on the altimetry, with a reverse anomaly along the coasts and mesoscale structures o↵shore, parallel to the coastline. Moreover, the surface circulation has an influence on the deeper water masses with the presence of highly saline water at 200 m, the depth of the Persian Gulf Water. During the winter monsoon these waters are found West of 56 o E, whereas in summer a front is found around 17 o N. This front can be due to the upwelled water In the next chapter, to complete this study on the marginal seas, we will focus on the Sea of Oman.

Chapter4. A seasonal dipolar eddy near Ra's Al Hamra (Sea of Oman) at Ra's Al Hadd, and stays along the coast, whereas in summer, it spreads eastward under the influence of the Ra's Al Hadd jet composed of a dipole with a cyclone North and an anticyclone South.

As in the previous section, we used ARGO floats to study the circulation in the Sea of Oman. An analysis reveals the presence of recurrent cyclonic motion in the western part of the basin and anticyclonic in the eastern part. As underlined in the previous chapter, these mesoscale eddies show a barotropic signature and thud influence the deeper water masses. Selecting 3 floats drifting in the Sea of Oman and in the northern Arabian Sea in spring, we study their trajectories under the influence of the dipole, and their measurements of water masses; we compare these trajectories to the bathymetry and to altimetry. Decomposing the time series of the altimetry into its principal mode of variability, via EOF (Empirical Orthogonal Function), we were able to understand the recurrence of this dipole. Furthermore to have a better understanding of the dynamic of this basin, on this inter monsoon, we use a quantitative approach, the modified thermohaline climatology, at the Persian Gulf Water mid depth to estimate preferential pathways in the basin. From the altimetry, and using ARGO constructed hydroglogy, we compute the geostrophic velocities, deformation and Okubo-Weiss fields at a specific date in order to study the influence at depth of the mesoscale structures. In a last part, we focus on the intensification of the anticyclone following the arrival of Rossby waves, emitted from the eastern Arabian Sea, and the seasonal reversal of the wind stress curl around the cape of Ra's Al Hadd. The results are presented in the following article published in 2013 in Ocean Dynamics.

Introduction

The Arabian/Persian Gulf (hereafter the Persian Gulf), located between the Arabian Peninsula and Iran, is a shallow sea, with maximum depth of 100 m near the Straits of Hormuz; the Persian Gulf is about 1,000 km long and 350 km wide, and is oriented in a NW-SE direction; it opens on the Sea of Oman (or Gulf of Oman) via the Straits of Hormuz which are about 110 m deep and 56 km wide. The Persian Gulf experiences fast tidal currents, especially near the straits and around islands, but residual currents are weak (Pous et al. 2012). The dominant northwesterly winds and the density gradients in the Persian Gulf induce a circulation that is mostly cyclonic, at least in its southeastern part, with an inflow of fresh Indian Ocean surface water from the Sea of Oman (Yao and Johns 2010a;Pous et al. 2013). The Persian Gulf is subject to strong evaporation (more than 1.5 m of freshwater evaporates per year) which dominates precipitation; therefore, salty waters are produced in its northern and southwestern regions. These two regions feed a salty outflow through the straits, with maximal respective contributions in summer and in winter (Reynolds 1993;Bower et al. 2000;[START_REF] Swift | Formation and circulation of dense water in the Persian Gulf[END_REF][START_REF] Yao | A hycom modeling study of the persian gulf : 2. formation and export of persian gulf water[END_REF]Pous et al. 2004a). The temperature and salinity of the outflow in the Straits of Hormuz are about 21 • C and 40.8 in winter and 25 • C and 39.3 in summer. On the 4.2. Abstract contrary, the Sea of Oman is a deep basin with a narrow shelf; maximal depths are in excess of 3,000 m. There, the winds are more variable and are influenced by the monsoon. The circulation is variable (mesoscale eddies grow and decay in this sea); upwellings can occur near the Iranian coast (Pous et al. 2004b).

Once in the Sea of Oman, the outflow from the Persian Gulf dives down the continental slope near of the Straits of Hormuz, mixes with surrounding waters, and adjusts hydrostatically at about 250-m depths along the coast of Oman. This outflow water is called Persian Gulf Water (PGW). The outflow transport is 0.12-0.15 Sv [START_REF] Johns | Observations of the seasonal exchange through the Straits of Hormuz and the inferred heat and freshwater budget of the Persian Gulf[END_REF][START_REF] Yao | A hycom modeling study of the persian gulf : 2. formation and export of persian gulf water[END_REF]. Available observations indicate that this outflow follows the Omani coast down to Ras al Hadd in summer and then bifurcates southeastward; in winter, the PGW outflow progresses beyond Ras al Hadd and follows the southern Omani coast (Bower et al. 2000;Pous et al. 2004b). The salinity of the PGW outflow decreases from 39.5 to 37.5 between the Straits of Hormuz and Ras al Hamra. The thermohaline properties of PGW along the coast of Oman show two strong decays, one in the Straits and one in the western part of the Sea of Oman (see the local T/S diagrams in Pous et al. (2004b)). Also, a few measurements identified isolated patches of PGW in the middle of the Sea of Oman. A single occurrence of a well characterized lens eddy of PGW is reported in [START_REF] Senjyu | High salinity lens from the Strait of Hormuz, in: Offshore Environment of the ROPME Sea Area after the War-Related Oil Spill[END_REF].

Recent modeling work suggests that the PGW outflow may leave the Omani coast near Ras al Hamra [START_REF] Ezam | Numerical simulation of spreading of the Persian Gulf outflow in the Oman Sea[END_REF], but these theoretical results await a clear observational confirmation. Unfortunately, oceanographic cruises spanning the whole Sea of Oman are rare. Therefore, Argo float data are used here. Previous work has already shown that the trajectory and thermohaline sections of these floats carry important information on the mesoscale dynamics in the Arabian Sea (Carton et al. 2012).

Our purpose is to confirm that PGW can be expelled from the outflow current near Ras al Hamra, to determine when and how this event takes place, and finally to explain the mechanism by which this occurs. After presenting the data and methods (Section 2), the analysis of Argo float trajectories and hydrological sections will show that PGW expulsion from the coast of Oman occurs preferentially in spring and occurred at least thrice in the past 5 years (Section 3). Then, we will present horizontal maps of temperature and salinity at the depth of the PGW obtained by interpolation of float data, and we will contrast the situations in spring and in autumn (Section 4). We will also show that a dipolar eddy exists and remains quasi-stationary near Ras al Hamra when PGW ejection occurs. Finally, we will relate the existence of this dipole with Rossby waves and the wind stress field at that period (Section 5), and conclusions will be drawn.

Material and methods

Argo float data were obtained from the Coriolis data center at IFREMER in Brest (France). Twenty-six floats drifted in the Northern Arabian Sea and in the Sea of Oman. Twelve of them covered the Sea of Oman only. The lifetime of these floats ranged from a few months to more than 4 years. Their cycle period (the time interval between two surfacings of the float) ranged from 4 to 10 days. Wrong hydrological data were discarded and so were floats which presented too many erroneous measurements. The accuracy of the float data was ±0.002 • C in temperature, ±0.005 in salinity, 1 m in depth, but data were provided only every 10 m from the surface down to 50 m depth, every 25 m down to 150 m, every 50 m down to 300 m and then every 100 m. Therefore, the water masses in the upper 300 m of the ocean were sampled over 13 vertical points. Here, three floats are presented in detail, to illustrate the expulsion of PGW from the coast into the central Sea of Oman or into the Arabian Sea, in recent years.

Each float provided many vertical profiles of temperature and salinity which are displayed here along the path of the float (this is called "temperature section" or "salinity section" below, though the abscissa is temporal). Note that the Argo floats are positioned only when they surface. Therefore, we call "float trajectory" the segments linking the successive float surfacings. The deep (and exact) float trajectory is unknown. The float trajectories will be segmented to display separate events.

We obtained the Generalized Digital Environment Model climatology from http://www.usgodae.org/ and used it to create horizontal maps of temperature and salinity; these maps were then modified by inserting the float data according to the following procedure. First, each float measurement was converted into an anomaly by substracting the climatological value for the given month and position. Second, this pointwise anomaly was extended spatially with an isotropic Gaussian correlation function, weighted by the local number of data (the radius of the Gaussian was chosen as L = 50 km, typical of a mesoscale eddy). This correlation function was normalized to unity where data lay. Third, the resulting anomaly was added to the climatology.

To determine the surface circulation, altimetric maps were obtained from the CNES/AVISO data center. These maps display absolute dynamic topography (ADT), with 1/4 × 1/4 degree resolution on a Mercator grid and daily interpolation; the mean sea level was a 7-year mean of the sea surface topography. The error on along-track measurements of sea level was 3-4 cm [START_REF] Fu | Satellite altimetry and Earth sciencesa handbook of techniques and applications[END_REF]. For the years 2008-2012, three altimetric satellites (Jason 1 and 2, and Envisat) were operational. For the Jason satellites, the inter-track distance at the Equator is 315 km and the repetitivity period of the orbit is 10 days; for Envisat, they are 80 km and 35 days. From the along-track data, gridded Chapter4. A seasonal dipolar eddy near Ra's Al Hamra (Sea of Oman) data were computed by AVISO via data merger and optimal interpolation (see (Ducet et al. 2000)). The quality of the maps strongly depends on the quality of the mean topography of the ocean surface, calculated from hydrographic and satellite data. More information on Mean Absolute Dynamic Topography (MADT) production is provided in (Gaspar and Ogor 1996;Labroue et al. 2006;[START_REF] Traon | Mesoscale mapping capabilities of multiple-satellite altimeter missions[END_REF]Le Traon et al. 2001).

To outline the cyclonic or anticyclonic motions at the sea surface, we created maps of "MADT anomaly," which is the difference between local MADT and its instantaneous spatial mean (note that this anomaly is different from SLA where the pointwise time average has been substracted from MADT).

These sea surface elevations were also studied from a statistical point of view to show their recurrence in time; empirical orthogonal functions (EOF) were calculated to determine the spatial structure of variability; seasonal time averages were also computed. The surface currents were computed via geostrophy, from the pressure anomaly associated with the sea surface elevation. The deep currents were computed from hydrology via the thermal wind equations, using the surface currents as a reference.

To assess the dynamical influence of the winds on the surface circulation, winds were retrieved from the CERSAT data center with 1/4 × 1/4 degree resolution and monthly mean series were obtained by collocation of scatterometer data. Also, high-frequency winds were obtained from the European Center for Medium-Range Weather Forecast model, at 1/8 × 1/8 degree resolution with 6-h time lag. The wind stress curl was computed from these latter data. For the monthly means, the curl was time-averaged and compared with that derived from the lower resolution data.

Selected trajectories and hydrological sections of Argo floats in the Sea of Oman

In the Argo database, three floats were selected which exhibited loops in the western Sea of Oman, and offshore motion from the coast of Oman or from the coast of Iran. These are Argo floats 2900804, 2901370, and 6900902. Figure 1 presents their respective domains of evolution over the regional bathymetry.

Analysis of Argo float 2900804

Float 2900804 was deployed December 14, 2007 near 24.51 • N, 59.15 • E. It was retrieved June 24, 2008 near 24.02 • N, 58.05 • E. It performed 4-day cycles with a parking depth at 1,000 m. Its trajectory shows many cyclonic loops in the center of the western Sea of Oman. Figure 2 details small loops with radii of about 40 km, northeast of Ras al Hamra from March 3 to April 16, 2008, and a larger loop along the continental slope from Iran to Oman, with a radius of 80 km, from May 22 to June 23, 2008. Such a large cyclonic loop of a float in the Gulf of Oman had already been described by Pous et al. (2004b).

These loops are superimposed on the altimetric maps averaged over the given period. During the first period, the surface flow was composed of a cyclone near 58 • E and of an anticyclone near 60 • E. The float trajectory looped mostly in the surface cyclone. In fact, the altimetric maps showed few changes in the surface circulation patterns from early February 2008 until end of April 2008, and this surface dipole was very persistent.

During the second period, a clear correlation also existed between the surface topography and the float trajectory. The surface geostrophic currents reached about 0.3 m/s, and the average Lagrangian velocity for the float during this period was 0.2 m/s in the same direction as the surface currents. The magnitude of this float velocity was comparable to that of currents at 900-m depths measured by Pous et al. (2004b). Indeed, the cyclonic and anticyclonic gyres in the Sea of Oman have a deep reaching dynamic signature. Note also that, in June 2008, a small anticyclone appeared near 58 -59 • E, between the large cyclone near 57 • E and the large anticyclone again near 60 • E.

In brief, from early February until end of June 2008, the surface circulation in the western Sea of Oman evolved very slowly, and the currents at 1,000-m depths (obtained from the float displacements) faithfully reflected the surface flow.

The salinity section along the whole float trajectory (not shown) indicates that two high-salinity events occurred between 200 and 400 m depths, one end of March, beginning of April and the other end of May, and in June; in both cases, the float recorded salinities above 37. Figure 3 zooms in on these two events.

On March 15, the salinity maximum at 200-m depths was measured at 24.5 • N, 59 • E, in the Sea of Oman, nearly on the axis between the cyclone and the anticyclone. The fresher water measured on March 23 indicates that the float also sampled the edge of the salinity anomaly measured a week earlier. A second salinity anomaly was measured, at 300-m depths, on March 27. This point was close to the continental slope and was likely to be located in the PGW slope current. Therefore, the salinity anomaly measured in mid Sea of Oman was probably a detached fragment from the slope current of PGW.

From May 22 to June 23, a very marked salinity anomaly (reaching 37.5) was measured by the float in the 200-300-m depth range (not shown). These recordings were performed by the float along the large cyclonic loop at the head of the Sea of Oman, firstly northward along the Iranian continental slope, then near the Straits of Hormuz, and finally southward along the Omani coast. Thus, PGW could be found all along the northern continental slope of the Sea of Oman at that time. Such spread out PGW signature could have been formed:

-either by an outflow which changed direction with time (e.g., firstly along the Iranian coast, then along the Omani coast), -or by an outflow with multiple branches, -or by short-term pulses of this outflow (Banse 1997), -or finally by the cyclonic recirculation of PGW in the western part of the Sea of Oman, starting from the ejection of PGW from the slope currents near Ras al Hamra.

The March-April event supports this latter interpretation. From April 2 to June 30, 2011, the float performed two cyclonic loops, with radii 60 km, east of Ras al Hamra, but with a superimposed westward motion (see Fig. 4). Then, from June 30 to December 14, 2011, the float drifted along the Iranian coast, first eastward and then westward. In a third stage (shown on Fig. 5), the float performed cyclonic loops near 24.5 • N, 58 • E, with radii 30-40 km, near the center of the Sea of Oman, from December 14, 2011 to April 24, 2012. Finally, a smaller-scale cyclonic trajectory with radii 25-30 km was followed by the float until mid June, west of 58 • E (see again Fig. 4).

Figure 5 shows two segments of the float trajectory, superimposed on the altimetric maps averaged over each period. In spring 2011, the float performed two counter clockwise loops correlated with a surface cyclone west of a large anticyclone, with a radius of 40 km. In April 2012, the float performed two cyclonic loops west of Ras al Hamra, mostly following the surface cyclone with the same radius as in spring. The correlation between the float trajectory and the altimetric signal was weaker here than in 2008; the radii Chapter4. A seasonal dipolar eddy near Ra's Al Hamra (Sea of Oman) The salinity section along the float trajectory shows the presence of PGW at 150-350-m depths, with salinity maxima above 37, in spring 2011 and in spring 2012 (see Fig. 6); this occurred, as the float was looping cyclonically west of Ras al Hamra. During both springs, the altimetric maps show the presence of the dipole at this location. In summary, the presence of a dipolar eddy, with a cyclone west of Ras al Hamra and an anticyclone east of this cape, most likely ejected salty PGW into the middle of the Sea of Oman; it also advected PGW across the head of the Sea of Oman thereafter. parking depth at 700 m. After an anticyclonic and a cyclonic loop near Ras al Hadd, it drifted northeastward towards the coast of Pakistan. There, it performed two small anticyclonic loops (from November 2011 until February 2012; see Fig. 7), followed by a large cyclonic loop (from February to May 2012, shown on Fig. 8) and then by an oscillatory motion at 24 • N between 63.5 • E and 64.5 • E (until June 2012). The part of its measurements of interest for the present study is that performed between November 2011 and May 2012. The large loop in the surface float trajectory, with radius close to 50 km, clearly followed the surface motions as shown by Fig. 8. The smaller loops, with radii smaller than, or equal to 30 km, were more poorly correlated with altimetry (these radii are below the resolution of the gridded altimetric data). The float velocity along these loops (and in particular along the large loop) was about 4 cm/s; this velocity is about half the estimated swirl velocity of the surface cyclone at 64.5

• E, 23 • N in March-May 2012.
The time evolution of sea surface height (not shown here) shows the formation of the small surface anticyclone end of January 2012, near 24.5 • N, 64 • E, surrounded by two small cyclones at 24.5 • N, 62.75 • E and at 24 • N, 65 • E. The small anticyclone was then advected southwestward by the small cyclone on its western side, hence the southwestward drift of float 6900902 during that period, superimposed on the loops. The small anticyclone reached 23 • N, 62 • E in early March 2012, while the two small cyclones merged on its northeastern side. The large cyclone in which the float was trapped in April was the result of this merger.

The time-versus-depth section of salinity is displayed on Fig. 9; contrary to the September 2011-January 2012 period, when salinity was lower than 36.3 at 200-400m depths (not shown here), the salinities in March-May were in excess of 36.5, especially in the 200-300-m depth range. Associated with temperatures around 19 • C, this water mass was clearly PGW. The presence of PGW at such eastern longitudes must be the result of its expulsion from the Omani continental slope earlier on (compatible with our previous observations), and of its advection along the Iranian and Pakistani coasts, either via currents or by eddies. The small cyclone-anticyclone pair could have advected this PGW away from the Pakistani coast.

In summary, these observations show expulsion of PGW from the slope current in April-June due to quasi-steady eddy pairs near Ras al Hamra. These eddies clearly had a deep-reaching flow field, as indicated by the float motions. This high-salinity water then recirculated westward (and cyclonically) in the Sea of Oman and eastward along the Iranian and Pakistani coasts, where it was again expelled offshore by eddies. To provide more general evidence of PGW expulsion from the Omani continental slope, seasonal salinity maps are now established via interpolation of float data, and recurrence of the dipole situation is searched in altimetric maps.

Hydrology

In Appendix 1, we show that the horizontal temperature (or salinity) fields obtained with our method or with Ocean Data View (ODV; [START_REF] Schlitzer | Ocean data view user's guide[END_REF])), coincide. Below, we show results obtained with our interpolation method.

Figure 10 compares the salinity at a 250-m depth in the Sea of Oman, in spring, for the climatology; for the modified climatology over years 2002-2012, the dipolar eddy was observed at the surface during the years 2008, 2011, and 2012.

The effect of using float data is to slightly reduce the value of salinity by 0.1 in the western end of the Sea of Oman (near 56.5 • E); but the relative error is maximum there due to a lack of float data (see Appendix 1). On the contrary, they increase salinity by 0.2 between 57.5 • E and 59.5 • E (a region where many profiles are available). A marked salinity gradient extends northeastward from Ras al Hamra in the modified climatology, whereas this region had a weak gradient in the original climatology. This is in agreement with the previous finding of recurrent PGW expulsion near Ras al Hamra in spring. Finally, the 36.4 salinity isoline bulges near 61 • E, 24 • N in the modified climatology, an indication of an eastward salt export, in agreement with the observations of float 6900902.

When only the 2008, 2011, and 2012 data are used, this effect is even more dramatic: salinity close to 36.9 now reach the Iranian coast near 58.5 • E, and a 36.6 salinity patch is found at 62 • E, 24 • N. This clearly illustrates the salt export associated with PGW expulsion at Ras al Hamra and its cyclonic and eastward recirculations.

Figure 11 presents similar maps, but for autumn. The inclusion of float measurements has an even stronger effect though few data are available again in the western end of the Sea of Oman (which limits the degree of confidence in results there). Several floats south of Ras al Hadd measured salinities in excess of 37.2. Since such salinities were not always present along the Omani coast up to the Straits of Hormuz, during this season, these strong anomalies must have been occasional and local bursts of salty water. Such bursts could be related to pulses in the outflow of PGW at Hormuz (Banse 1997;[START_REF] Johns | Observations of the seasonal exchange through the Straits of Hormuz and the inferred heat and freshwater budget of the Persian Gulf[END_REF]Thoppil and Hogan 2009;[START_REF] Bidokhti | The structure of the Persian Gulf outflow subjected to density variations[END_REF] but the concentration of this salty anomaly near Ras al Hadd more likely indicates an instability of the PGW slope current at that location in autumn.

Sea surface elevation, surface and deep currents

Maps of sea surface elevation (from satellite altimetry) show that the dipolar eddy was present and stationary for several months near Ras al Hamra in 1994Hamra in , 1995Hamra in , 1997Hamra in , 1998Hamra in , 2005Hamra in , 2007Hamra in , 2008Hamra in , 2011Hamra in , and 2012, and , and In Appendix 1, we show that the relative error in sea surface elevation is moderate at the location of the dipolar eddy.

On Fig. 12, we present the geostrophic currents at 300-m depths computed via thermal wind balance, with the hydrological data and using the surface current as a reference. This surface current itself was calculated from the sea surface elevation given by satellite altimetry of March 12, 2008. The currents at this depth closely followed the surface cyclones and anticyclones (see for instance Fig. 2). The strongest currents were located between the two eddies near 59 • E, 24.5 • N. The flow converged towards this point from Ras al Hamra and then diverged towards the Iranian coast. Figure 12 also shows the deformation field at 300-m depths. Deformation σ was computed as follows:

σ 1 = ∂ x v + ∂ y u, σ 2 = ∂ x u -∂ y v, σ 2 = σ 2 1 + σ 2 2
We can clearly see that the float followed the regions of maximum deformation. This supports our interpretation that the surface eddy pair exerted a strong deformation on the PGW slope current. This deformation pulled some PGW out of the slope into the deeper part of the Sea of Oman.

The Okubo-Weiss theory [START_REF] Okubo | Horizontal dispersion of floatable particles in the vicinity of velocity singularities such as convergences[END_REF][START_REF] Weiss | The dynamics of enstrophy transfer in two-dimensional hydrodynamics[END_REF]) states that a Lagrangian conserved quantity has a spatial gradient increasing or decreasing with time, depending on the difference between deformation and relative vorticity. When relative vorticity dominates deformation, the gradient diminishes (and conversely). Therefore we have calculated the Okubo-Weiss quantity σ 2ω 2 (where ω = ∂ x v -∂ y u is relative vorticity), from the horizontal velocity field at a 300-m depth.

This quantity is plotted in Fig. 12 for March 12, 2008, also. Along 24.5 • N, two regions of negative Okubo-Weiss quantity correspond to the two eddies where any materially conserved tracer would recirculate and be trapped. On the contrary, the two regions of positive Okubo-Weiss quantity, along 59 • E correspond to strong deformation which can pull fragments of PGW away from the slope currents. Note also that this quadrupolar structure is very characteristic of the deformation field created by an eddy dipole (see for instance Carton et al. (2002) for another oceanic example).

Figure 13 presents the time average, and the EOFs, of mean ADT over spring 2008, 2011, and 2012. Clearly, the time average of MADT indicates an eastward increase of the sea surface height. If the spatial mean is substracted from MADT, the geostrophic circulation is cyclonic west of 58 • E and anticyclonic west of 59 • W . This dipolar structure is robust in a seasonal average repeated over 2008, 2011, and 2012. The EOFs were also computed from the time series of mean ADT over spring of these 3 years. The first EOF shows an anticyclonic anomaly near Ras al Hadd and represents 73 % of the variability during this period. It has a spatial structure similar to that of the time-averaged MADT. The second EOF, which represents 9.5 % of the variability, also has a dipolar structure, but closer to Ras al Hamra, with a cyclonic region east of 58.5 • E and an anticyclonic region between 59 • E and 61 • E. Finally, the third EOF, with 5.6 % significance, has a cyclone near 57.5 • E and an anticyclone near 59 • E. In summary, the first and third EOF reproduce the spatial structure of the dipole while the second EOF reproduces an opposite signed dipole, which is likely to represent the first-order fluctuations of the seasonal structure around its mean.

From the CNES/AVISO Live Access Server, one can obtain the time evolution of surface eddies in the Gulf of Oman, for every year when the dipole appeared, and also for the year before. This time series is shown in Appendix 2 (see Figs. 18 and19). The evolution is always similar: in September of the year before each spring dipole, the circulation in the Gulf of Oman is dominantly cyclonic. In winter, an anticyclonic circulation appears at the mouth of the Gulf. This anticyclonic patch moves westward until it reaches Ras al Hamra in February-March. This evo-Chapter4. A seasonal dipolar eddy near Ra's Al Hamra (Sea of Oman) lution suggests that a westward traveling wave could be responsible, in part or totally, for the development of this anticyclonic patch. The dipolar situation is then established until June or July. In July or August, the anticyclone weakens and the circulation returns to cyclonic in autumn and early winter.

To investigate the existence of such a westward wave, we present a Hovmoller diagram of the MADT anomaly at 24.5 • N from 2000 to 2012 (see Fig. 14). The sea surface height evolution shows westward propagating anomalies. It also indicates an alternation between positive and negative anomalies with an annual cycle. Similar wave propagation across the Indian Ocean has already been shown to be long baroclinic Rossby waves with annual period [START_REF] Brandt | Annual Rossby waves in the Arabian Sea from TOPEX/POSEIDON altimeter and in situ data[END_REF]. For Rossby waves, the phase speed is This leads to a phase speed c = 2.5 cm/s in agreement with the observations. Thus a long Rossby wave on the first baroclinic mode is an appropriate mechanism to explain the westward traveling sea surface anomaly along 24.5 • N. Figure 14 also shows an intensification of the sea surface height anomalies near 61 • E between December and March. As in [START_REF] Brandt | Annual Rossby waves in the Arabian Sea from TOPEX/POSEIDON altimeter and in situ data[END_REF], this intensification may be due to the wind stress near 60 • E in winter. We will now investigate the wind stress to support this hypothesis.

c = ω k = - β k 2 + 1/R 2

Conclusions

In summary, Argo float trajectories, and temperature and salinity recordings, in the Sea of Oman, have shown that the currents at 200-300-m depths could be directed offshore in spring near Ras al Hamra, and that fragments of warm and salty Persian Gulf Water could then be stripped away from the slope current of PGW. The analysis of altimetric maps from the CNES/AVISO center has shown that a mesoscale dipolar eddy was often present in spring near Ras al Hamra, with a cyclone west of this cape and an anticyclone east of it. This orientation leads to a flow between the eddies that is offshore. These eddies are deep reaching dynamically, as was evidenced by the trajectory of the Argo floats. This dipole could have been formed by the autumn cyclonic circulation joined with anticyclonic vorticity production at the mouth of the Sea of Oman in winter.

In more details, the paper presented the following results:

-Though a few previous papers have mentioned the possible detachment of PGW from the coast, the traditional view was that the PGW outflow hugs the Omani coast and veers southeastward at Ras al Hadd in summer and southwestward in winter; here, we have shown that PGW could be found both under the form of fragments in the middle of the Sea of Oman and as patches near the Iranian coast. This new picture results from the individual analysis of Argo float recording and from the construction of a new seasonal climatology including these data. These warm and salty anomalies in the middle or north of the Sea of Oman are all the more pronounced as only the years where the dipole was present were included in the climatology. -Furthermore, it appears that PGW does not only recirculate in the western Sea of Oman, but that it can also extend eastward along the Iranian and Pakistani coasts.

There, other eddying motions can extract fragments from this coastal patch and advect them offshore, as float 6900902 showed. The complex pattern of PGW extension due to turbulent motions will certainly be -The trajectory of the Argo floats often followed loops of different sizes, but most frequently measured radii varied between 30 and 60 km. The floats followed the surface currents, with an amplitude ratio of about 2/3 in velocity, especially in the presence of surface eddies. These eddies have a diameter varying between 150 and 250 km; often cyclonic in autumn in the Sea of Oman while anticyclonic eddies appear in late winter. A recurrent feature is the presence of the dipolar eddy near Ras al Hamra in spring. -The analysis of the circulation at 200-300-m depths (PGW depth) showed that the dipole creates a strong deformation along its axis, which is likely to be the cause of PGW extraction from the coast. This also appears in the structure of the Okubo-Weiss quantity. -The dipolar eddy observed in spring of several years seems to result from long baroclinic Rossby waves propagating at the mouth of the Sea of Oman; it could have been amplified by the wind stress near 60 • E during winter. It was noted that dipole formation occurred during the years when the positive SSH intensified near the coast of Oman, in the Rossby wave; this supports our interpretation. Another mechanism which could also contribute to the formation of this dipole is the interaction of the PGW outflow with bottom topography, leading to the instability of this outflow current. Indeed, in winter, the PGW outflow is deeper [START_REF] Bidokhti | The structure of the Persian Gulf outflow subjected to density variations[END_REF], and a baroclinic instability process (Pous et al. 2004b) triggered by topographic irregularities along the continental slope could lead to eddy detachment. Nevertheless, the specific correlation of this latter mechanism with the years of dipole formation still has to be established.

Two interesting extensions of this work may be the following:

1. Use a regional numerical model to obtain more insight in the three-dimensional dynamics of the Sea of Oman,

Conclusions

both for the effect of mesoscale eddies on the PGW outflow and for the generation and evolution of these mesoscale eddies; 2. Use other data (or again a numerical model), to provide a more complete view of the extent of PGW due to mesoscale variability.

Further work and complementary data

Another application of the EOFs on the altimetry, this time not on the ADT, but on the ADT anomaly, provided a better decomposition of the variability (see figure 4.1). As well, with the ANDRO database from IFREMER we were able to retrieve numerous ARGO floats from the Data Assembly Centers (DACs) in the region. Since this article the number of launched floats increased. The Sea of Oman has a particularly good coverage with about 500 measurements per month that allowed us to have finer temporal and spatial resolutions (see figure 4.2).

The first and second EOFs, with a representativity of 27% and 13% underline the presence of the anticyclone in the mouth of the Sea of Oman, for each year in March (see figure 4.1). As well, large gyres with radii of about 150-200 km and with alternating rotation directions are observed West and South of the anticyclone. The third EOF (9.2%) shows numerous structures of smaller sizes, about 100 km and lower intensity. Moreover, it points out the presence of eddies in the Sea of Oman, with the cyclone of the dipole about half of the observed years. The first two EOFs indicate a first mode of variation whereas the third EOF indicates a second mode, thus smaller structures, erased from the averaged altimetric anomalies.

At 200 meter depth, the mean depth of the Persian Gulf Water outflow, the ARGO floats (see figure 4.2, up) point out the ejection of the most saline and warm water East o↵ the Sea of Oman at 24 o N. This also corresponds to the position of a front in Sea Surface Temperature between the surface colder water in the Sea of Oman, and the warmer water in the Arabian Sea. This front, and the circulation of the Persian Gulf Water, match the position of the anticyclone at the mouth of the Sea of Oman.

Inside the basin, the most saline waters are found along the southwestern coasts and an ejection is observed o↵ Ra's Al Hamra. This correlates with the Sea Surface Temperature map, where the surface colder water are advected North at Ra's Al Hamra. As well, a cyclonic motion of the water mass is observed in the western part. This emphasizes the role of the mesoscale eddies on the pathway of Persian Gulf Water out of the Sea of Oman from a quantitative and climatological side. Furthermore the outflow is not confined to the coast of Oman as during the monsoons.

A study of the annual cycle of the Persian Gulf Water outflow is needed to fully understand the dynamics in the Sea of Oman. In the next chapter we study the circulation and seasonal variations in the Arabian Sea, with an emphasis on the Sea of Oman. We also focus on the characteristics of the Persian Gulf Water and on its interactions with the mesoscale, seasonal, eddies in the whole western Arabian Sea and process study through numerical model.

N o N o N o N o N o N o N o N
Third EOF (9.2%) for March 1992 to 2014 and amplitude Chapter5. Mesoscale variability in the Arabian Sea from HYCOM model results and observations : impact on the Persian Gulf Water path

Motivations for the research

The previous chapters allowed us to study first from a qualitative, then from a quantitative point of view the circulation in the Gulf of Aden, Arabian Sea and Sea of Oman, revealed by ARGO floats. The presence of mesoscale eddies, with a strong variability depending on the seasons and on the surface forcing has been emphasized. It has been observed that these eddies, or dipoles, have an influence on the water masses and the highly saline Persian Gulf Water outflow.

In this chapter we describe an averaged annual cycle in the northern Arabian Sea and Sea of Oman, with a focus on each season and from several points of interest. Firstly the atmospheric forcing, net heat flux and winds are studied with their impact on the ocean, inducing variations of the sea elevation, the evaporation, the creation of water masses. Secondly, a strong focus is put on the mesoscale eddies, from their formation to their evolutions and interactions with the surface forcing and with other structures, as eddies or Rossby waves. Thirdly, we concentrate on the circulation of the Persian Gulf Water over an average year, with the variations of its outflow and characteristics.

This study is achieved using tools described earlier for the observational data, surface fields from satellite dataset and reanalysis (winds, altimetry and heat fluxes). These averaged observations are compared to the outputs from the HYCOM model on this region, in order to validate it. Fields of velocities and of thermohaline characteristics are obtained with high spatial and temporal resolutions, thus allowing us to follow the di↵erent structures along their life cycles and the related mechanisms. The results are provided in an article submitted to Ocean Science in January 2015.

Abstract. The Arabian Sea and Sea of Oman circulation and water masses, subject to monsoon forcing, reveal a strong seasonal variability and intense mesoscale features. We describe and analyze this variability and these features, using both meteorological data (from ECMWF reanalyses), in situ observations (from the ARGO float program and the GDEM -Generalized Digital Environmental mode -climatology), satellite altimetry (from AVISO) and a regional simulation with a primitive equation model (HYCOM -the Hybrid Coordinate Ocean Model). The model and observations display comparable variability, and the model is then used to analyze the three-dimensional structure of eddies and water masses with higher temporal and spatial resolutions than the available observations. The mesoscale features are highly seasonal, with the formation of coastal currents, destabilizing into eddies, or the radiation of Rossby waves from the Indian coast. The mesoscale eddies have a deep dynamical influence and strongly drive the water masses at depth. In particular, in the Sea of Oman, the Persian Gulf Water presents several offshore ejection sites and a complex recirculation, depending on the mesoscale eddies. The associated mechanisms range from coastal ejection via dipoles, alongshore pulses due to a cyclonic eddy, to the formation of lee eddies downstream of Ra's Al Hamra. This water mass is also captured inside the eddies via several mechanisms, keeping high thermohaline characteristics in the Arabian Sea. The variations of the outflow characteristics near the Strait of Hormuz are compared with variations downstream.

Introduction

The circulation and water masses in the northern Arabian Sea and the Sea, or Gulf, of Oman are investigated here using a numerical ocean model at mesoscale resolution and in situ and remote sensing data. The focus lies here on the seasonal structure and variability of intense mesoscale features and on their impact on the outflow of Persian Gulf Water (hereafter PGW).

In the Indian Ocean, atmospheric and oceanic processes strongly reflect the monsoonal cycle. This study focuses on the northwestern part of this ocean, the Arabian Sea, north of 18 N. Figure 1 presents the basins of interest and geographical locations. In the Arabian Sea, the winter monsoon lasts from November to February, with steady northeasterly winds up to 10 m s 1 on average. The winds reverse in summer to blow from the southwest from June to September, with a stronger intensity, up to 14 m s 1 on average. The monsoons are interleaved by short transition periods in spring and fall when winds are weaker and variable in direction (see [START_REF] Defant | Physical Oceanography[END_REF].

These steady winds have several impacts on the ocean (see [START_REF] Schott | The monsoonal circulation of the Indian Ocean[END_REF]. They induce an Ekman transport during the monsoon seasons, chasing away the water from the coasts in summer and pushing towards them towards the coast in winter. Consequently this creates upwellings and downwellings along the coasts (see Shi et al., 2000). The most noticeable upwellings occur near the Somalia and Oman coasts, with vertical velocities up to 3 ⇥ 10 2

Published by Copernicus Publications on behalf of the European Geosciences Union. and 10 2 mm s 1 , during the summer monsoon with the production of very long cold filaments protruding from the coast with a transport from 4 to 8 Sv. Downwellings, mainly along the northeastern coast of the Arabian Sea, along India and Pakistan, have vertical velocities of comparable magnitude, up to 1.5 ⇥ 10 2 mm s 1 .

Abstract

The surface circulation is also constrained by the seasonal winds, particularly the coastal currents, changing direction with the monsoon. Along the western boundary of the Arabian Sea, the Oman Coastal Current (OCC) flows southwestward in winter and reverses in summer to the northeast. At the Ra's Al Hadd cape, it forms, in summer, an eastwardpropagating jet known as the Ra's Al Hadd jet (see [START_REF] Flagg | Upper ocean currents in the northern Arabian Sea from shipboard ADCP measurements collected during the 1994-1996 US JGOFS and ONR programs[END_REF], the most intense feature of the northern Arabian Sea, with velocities of 1 m s 1 and an associated transport of 2 to 8 Sv. Along the coast of India, the West Indian Coastal Current (WICC) flows northward during the winter monsoon, with decreasing intensity from 15 to 24 N (see Shetye et al., 1991), and southward in summer. Another noticeable feature of the western Indian coast is the presence of coastal Kelvin waves propagating from the Bay of Bengal during the winter monsoon (see [START_REF] Rao | Interannual subsurface variability in the tropical Indian Ocean with a special emphasis on the Indian Ocean dipole[END_REF].

Offshore, in the interior of the Arabian Sea, the circulation is dominated by mesoscale eddies (see Schott and Mc-Creary Jr., 2001) with sizes of 100-200 km. Some features are recurrent, such as the Ra's Al Hadd jet, composed of a dipole, or also an anticyclone off Ra's Ash Sharbatat, described by [START_REF] Flagg | Upper ocean currents in the northern Arabian Sea from shipboard ADCP measurements collected during the 1994-1996 US JGOFS and ONR programs[END_REF]. The bi-monthly averaged maps of Fig. 2 show the MADT (mean absolute dynamic topography; see the data and methods section for computation) anomaly. Positive anomalies correspond to anticyclonic flows and negative anomalies to cyclonic ones. These maps indicate the presence of recurrent structures from one year to another. The evolution of the MADT over the years is highly periodic and is correlated with the wind stress (see Fig. 3). In the Indian Ocean, variability is observed in a semiannual mode, but also at 135, 120 and 90 days, pointing out the different regime for each season, monsoons or intermonsoons. These spectral peaks are also observed in other fields in this region such as the sea surface temperature. Other periods, at 90, 75 and 60 days, were studied by [START_REF] Fu | Intraseasonal variability of the equatorial Indian Ocean observed from sea surface height, wind, and temperature data[END_REF].

In the Sea of Oman, the circulation is mainly composed of a western boundary current, intense in summer (up to 0.8 m s 1 ), protruding in the Ra's Al Hadd jet, and weak in winter (see [START_REF] Flagg | Upper ocean currents in the northern Arabian Sea from shipboard ADCP measurements collected during the 1994-1996 US JGOFS and ONR programs[END_REF]. The fall intermonsoon and early winter circulations were studied by Pous et al. (2004b), who observed the boundary current veering around Ra's Al Hadd, flowing southwestward and feeding the OCC. Evidence has been shown of seasonal or higher-frequency variability of the surface circulation in the Sea of Oman, which has been related to atmospheric forcing (seasonal wind stress or the effect of atmospheric cyclones) or to Rossby waves emitted from the eastern boundary (Pakistan and India; see [START_REF] Schott | Indian Ocean circulation and climate variability[END_REF]. In particular, mesoscale oceanic dipoles (cyclone + anticyclone), as well as isolated eddies, have been observed in the vicinity of Ra's al Hamra (see L 'Hegaret et al., 2013) and off Ra's al Hadd (see [START_REF] Wang | High salinity events in the northern Arabian Sea and Sea of Oman[END_REF].

Another impact of the seasonally steady winds, combined with the net heat flux of the tropical basin, is the high annual evaporation rate. The net evaporation-precipitation budget is large and positive in the Persian Gulf and in the Red Sea. This leads to the generation of water masses with very high salinities (above 40 psu) in these two basins (see Yao andJohns, 2010a, b, andBower et al., 2000). The present article focuses on the saline water flowing out of the Persian Gulf.

The Persian Gulf is a very shallow sea with an average depth of about 35 m and maximum depths around 100 m. Very salty and dense water, Persian Gulf Water (PGW hereafter), is formed via evaporation near the western boundary of the Persian Gulf (see [START_REF] Pous | A model of the general circulation in the Persian Gulf and in the Strait of Hormuz: Intraseasonal to interannual variability[END_REF]. This water flows through the Strait of Hormuz and spreads in the Sea of Oman and in the Arabian Sea (see Bower et al., 2000, and Fig. 1 for geographical locations and bathymetry). Observations show that the annual mean outflow transport of PGW through the strait is about 0.25 Sv, with little seasonal variation in transport (see [START_REF] Matsuyama | Vertical structure of a current and density front in the Strait of Hormuz[END_REF]Johns et al., 2000). But this transport can be affected by short-term pulses related to the effect of oceanic cyclones in the Persian Gulf (see Thoppil and Hogan, 2009). Outflowing PGW is colder and saltier in winter (Pous et al., 2004a).

Ocean

In the northern part of the Sea of Oman, PGW reaches an equilibrium depth between 150 and 250 m with a density ✓ between 26.3 and 27 (see Fig. 4). The PGW plume tends to deepen away from the straits. It can be found at 250-350 m depth south of the Arabian Peninsula (see Bower et al., 2000, Pous et al., 2004b, and Carton et al., 2012). Irregularity in the outflow transport and the presence of mesoscale eddies, as well as topographic accidents, can disrupt this deep current, which then sheds fragments of PGW (see L 'Hegaret et al., 2013). PGW can form patches and recirculate in the Sea of Oman. More prominently, an intrathermocline eddy of PGW was once observed near the straits, extending towards the middle of the Sea of Oman (see [START_REF] Senjyu | High salinity lens from the Strait of Hormuz, in: Offshore Environment of the ROPME Sea Area after the War-Related Oil Spill[END_REF].

In the Arabian Sea, the surface layer is occupied by the Arabian Sea High Salinity Water (ASHSW) down to 100 m depth, presenting high salinity and temperature due to the intense evaporation and low precipitation over the basin, with a ✓ around 23. River inflows are observed in late summer in the western Arabian Sea as a thin surface layer exiting the Indian coast. Below the surface layer, Indian Ocean Central Water is observed with temperature between 15 and 20 C, and with salinity below 36 psu; thus, it has a lighter density than PGW, with ✓ between 25 and 26. The deep and bottom layers are filled by Indian Ocean Deep Water and Antarctic Bottom Water, with ✓ under 27 and 27.5, respectively.

The objective of this paper is twofold. The first objective is to investigate and rationalize the mesoscale variability in the Sea of Oman and Arabian Sea, to characterize the different features, their three-dimensional structures and characteristics, and their seasonal variability. Then the study focuses on the mutual interactions of these features and their interactions with topography and atmospheric fluxes. Second, once the dynamical context has been studied, the characteristics of the PGW are investigated, as well as their seasonality and the various mechanisms of detachment of PGW fragments from the plume. To achieve these purposes, a primitive equation model is implemented regionally and the results of the simulations are analyzed in conjunction with synoptic in situ and surface data. The paper is organized as follows: the model characteristics and the origin and nature of data are first specified; the regional structure and seasonal variability of the water masses and currents in the region are then described from the two independent data sources; the mesoscale structures and their interactions are characterized and, finally, the impact of these structures on the PGW outflow is quantified.

Data, model and methods

The numerical simulations were performed with the Hybrid Coordinate Ocean Model (HYCOM) (see Chassignet et al., 2007). HYCOM is a primitive equation, hybrid coordinate model (isopycnal and z). The model variables are discretized on a horizontal Arakawa C-grid. The horizontal resolution is here 5 km and the model has 40 vertical levels. The model simulates both of the basins of evaporation, the Red Sea and the Persian Gulf, and extends to the south of India (see Fig. 1). The simulation is initialized with a state given by a lower-resolution global model (MERCATOR PSY3, 1/4 ) and is forced by realistic atmospheric fields (from the French Met Office: METEO FRANCE) cycling over 2011. The open boundaries, south and west, south of India, are also forced in salinity, temperature and currents by MERCATOR. The tidal forcing is provided by Topex and is included as a boundary condition. The vertical mixing is a K-profile parameterization (KPP; see [START_REF] Large | Oceanic vertical mixing: A review and a model with a nonlocal boundary layer parameterization[END_REF]. The model provides a PGW slightly lighter with ✓ = 26 instead of 26.3 due to higher temperature (see Fig. 4). Lighter water from river discharge is poorly sampled by the ARGO floats that most often drift at 1000 m depth. Indeed, this water mass is essentially confined above the shallow shelves of the western India coast. Nevertheless, the water masses show no other difference between HYCOM and ARGO profiles. An 11-year simulation starting in 2011 was produced. The first 5 years correspond to the spin-up of the model and were discarded. The spin-up was considered achieved when the mesoscale circulation was well established and the properties of the main water masses reached their typical values. Only the last 6 years of the simulation were analyzed; in our timeline, they are denoted 6 to 11.

To study the thermohaline variability over the seasons, we used a modified GDEM (generalized digital environmental model) climatology. GDEM is a monthly climatology constructed from MOODS (Master Oceanographic Observational Data Set) profiles, with 72 vertical levels and a 1/4 ⇥ 1/4 horizontal resolution (see [START_REF] Carnes | Description and Evaluation of GDEM-V 3.0[END_REF], for climatology construction). This climatology provided a good description of the Arabian Sea and the bordering gulfs and seas, particularly the Sea of Oman. To improve the horizontal resolution and to follow the different water masses, we added profiles from ARGO floats. These long-lived drifters provide high-quality profiles for temperature and salinity. We used the ANDRO data set for floats released before 2010 and the CORIOLIS (France) and INCOIS (India) databases for later ones. Over 100 floats were found to be drifting north of 18 N in the Arabian Sea between 2002 and 2014, which corresponds to 550 to 600 profiles per month, on average.

The surface fluxes that were used for analysis were issued from a monthly meteorological model reanalysis by ECMWF. This reanalysis provided monthly climatologies of wind, heat, evaporation and precipitation fluxes over 40 years with a resolution of 0.7 . The horizontal resolution is low but sufficient to resolve the mesoscale at these latitudes, and this provided a data set on a long period of time, necessary to compute a climatology. We also used altimetric data from AVISO over 22 years, from 1992 to 2014. From the MADT, the instantaneous spatial average was subtracted to obtain MADT anomalies.

Different parameters are calculated from the results of the HYCOM simulation. The deformation is the sum of the squared strain and shear, = 2 strain + 2 shear with strain = @V @x + @U @y , and shear = @U @x @V @y the shear. The anomaly of Ertel potential vorticity (PVE) is computed from the total PVE PVE = 1

⇢ 0 . ✓ f + @V @x @U @y . @⇢ @z @V @z . @⇢ @x + @U @z . @⇢ @y ◆ (1)
from which we substract a PVE value calculated out of the eddies and jets.

Ocean

Mesoscale structures, formations and evolutions

Three mesoscale structures are presented hereafter: coastal currents, eddies and waves. Their formations, structures, evolutions and seasonality are studied in the Arabian Sea and in the Sea of Oman.

Coastal currents

The first type of mesoscale structure studied here is the coastal current. The monsoonal winds induce a net Ekman transport offshore in summer and shorewards in winter along the northern and western coasts of the basin. This leads, respectively, to a lowering (with the upwellings) or to an uplifting of the sea surface in these regions, with the formation of geostrophically balanced coastal currents. A contrast is observed between the coasts and the interior of the basin in Fig. 2. From May to September, the MADT anomaly is negative along the coast, while it is positive in the center of the basin; the situation reverses in November until January.

This coastal current is about 150 km wide, with an associated velocity up to 0.7 m s 1 near the strongest eddies and about 0.3-0.4 m s 1 elsewhere, directed southwest in Fig. 5 (center), in the HYCOM simulation during the winter monsoon. Offshore, northeastward velocities are observed, and the map of the Ertel potential vorticity anomaly indicates a change of sign between the offshore and coastal values. The likeliness of the barotropic instability can be estimated for this structure by evaluating the Rayleigh-Kuo criterion, which states that @ 2 U h @y 2 changes sign in the region if the flow is unstable; U h is the horizontal velocity. Here, taking into account the slope of the current, U h /L 2 and p 2/2 have the same order of magnitude of 1.3 ⇥ 10 11 m 1 s 1 , with U h = 0.3 m s 1 , L = 150 km and = 2 ⇥ 10 11 m 1 s 1 . Thus, barotropic instability can occur for the coastal current.

The cross section in Fig. 6 shows that, in the upper few hundred meters, an inversion of the Ertel potential vorticity anomaly occurs vertically. The possibility of baroclinic instability can be assessed for this current using a simple two- layer Phillips model. With two equally thick layers and opposite surface and deep flows (see [START_REF] Vallis | Atmospheric and oceanic fluid dynamics: fundamentals and large-scale circulation[END_REF], the condition for existence of baroclinic instability is that the vertical shear of horizontal velocity U s be larger than R 2 d /4. Here R d ⇡ 40 km, so that the threshold for velocity shear is about 0.01 m s 1 . Thus the possibility of baroclinic instability for this coastal current is validated. Maximum baroclinic instability is obtained under such conditions for = 2⇡R d /1.8 ⇡ 140 km (theoretical value for a large shear), which is comparable with the largest western eddy diameters in observations (approximately 200 km). Note that a similar estimate is obtained with the Eady model, ⇡ 4R d , corresponding to a most unstable wavelength of 160 km (see also [START_REF] Smith | The geography of linear baroclinic instability in Earth's oceans[END_REF].

Finally, the growth rate of such unstable waves is = 0.3U s /R d in the Phillips model, or about 1/5 days, with U s = 0.3 m s 1 . The Eady model gives the same expression and result. On MADT maps (see Fig. 7), eddies are fully formed within 20 to 25 days after the appearance of the me-anders. This duration is about 4 1 (or four linear periods of instability), which takes into account the time necessary for wave growth and nonlinear interactions. Such a duration was commonly observed for eddy formation and detachment from an unstable jet in previous studies of baroclinic instability (see [START_REF] Flierl | Vortex formation by unstable oceanic jets[END_REF]. This instability of the coastal jets leads to the second type of mesoscale feature in the basin, the mesoscale eddies.

Mesoscale eddies

Mesoscale eddies are the most intense dynamical structures in the northern Arabian Sea. They are observed mainly west of 65 E with increasing diameters from east to west, from 100 to 200 km (see Figs. 2 and8). The strongest eddies are observed along the coasts or aligned along the coastal currents. North of 15 E and south of Ra's Al Hadd, four intense eddies are found along the coasts; two to four eddies are ob- served in the Sea of Oman. These structures have a seasonal recurrence, as seen in Fig. 2, with dipoles forming off Ra's Al Hadd in summer, or off Ra's Al Hamra, in the Sea of Oman, in spring.

As seen in the previous section, their formation is linked to the destabilization of the coastal current, with the formation of cyclonic eddies in summer (anticyclonic in winter) along first the Iranian coast, then the Omani coast. The amplification of these eddies can result from a localized, steady and strong wind stress curl forcing. This occurs in particular in the Sea of Oman and along the western coast of the basin, for all seasons, and furthermore along the Indian coast in summer. The wind-induced relative vorticity is expressed as [START_REF] Dong | Island wakes in deep water[END_REF]. The strongest values are found near these coasts and in the Sea of Oman with an order of magnitude of 2.5 ⇥ 10 6 s 1 , positive in summer, negative the rest of the year. Comparatively this is about 35 % of the relative vorticity calculated from the sea surface height, with 7⇥10 6 s 1 for the mesoscale eddies. As shown in L' Hegaret et al. (2013), in spring to early summer near Ra's al Hadd, the wind veers clockwise at this period and imparts anticyclonic vorticity to the ocean, thus reinforcing an anticyclonic eddy. A further amplification of this vorticity was then interpreted by the same authors as the result of the interaction with an incoming upwelling Rossby wave; this process will be studied in more detail here.

1⇠ = ! k . ! r ^ ! ⌧ ⇢ 0 H 1t (see
The eddies are intensified in the upper water column, with velocities up to 0.7 m s 1 from the surface down to 150 m depth (see Fig. 9). The velocity section shows the vertical influence of those eddies. From the HYCOM simulation, correlation between the surface and at-depth velocities can be estimated. Figure 10 (up) shows the correlation coefficient between the surface and 600 m depth velocities. Strong correlations are found around the eddies, with the exception of their centers, where tilting can slightly modify their axes. Zooming in on a specific eddy, the attenuation coefficient can be retrieved; thus, 70 % of the surface velocity intensities are recovered at 300 m depth, 50 % at 500 m, and up to 30 % at 1000 m. This vertical influence was also seen in the strong correlation between float drifts at 700 m and the surface geostrophic velocities (see L 'Hegaret et al., 2013). These mesoscale eddies induce an intense deformation field associated with their high and sheared velocities, having thus an impact on the water masses at depth.

Under the anticyclone and the cyclone, the vertical spacing between isopycnals shrinks and widens successively with a stronger deviation in the first 300 m (see Pous et al., 2004b). These isopycnal deviations have an influence on the water masses. The surface water, ASHSW, is sampled in the first 150 m; below, the IOCW (Indian Ocean Central Water) is colder and fresher with salinity under 36 psu, and between 200 and 300 m the PGW is observed. The core of the IOCW is thinner in HYCOM than in the in situ observations because of the depth of the PGW, which is found deeper than in the model. The eddies can deepen or elevate the depth of the PGW. Between the eddies, layers of ASHSW and IOCW wrap around the cyclone, injecting different water masses inside the eddy. The stratification is maximum at the depth of the PGW core and near the injection around the eddies.

Rossby waves

Time vs. longitude diagrams of the MADT anomaly from AVISO (see Fig. 11) reveal westward propagative anomalies from the Indian coast. Some of these anomalies have been identified as Rossby waves (see L 'Hegaret et al., 2013) with a wave speed of c = R 2 d of 21 cm s 1 and with a wavelength of about 210 km. These waves have also been observed more southwards by [START_REF] Heffner | Indian Ocean Rossby waves detected in HYCOM sea surface salinity[END_REF]. The Hovmoller diagrams from both AVISO and HYCOM point out the strong reversal of positive and negative anomalies along the eastern side of the basin followed by wave emission. Penetration of Kelvin waves along the coast of India from the Bay of Bengal is known to emit Rossby waves in the southern part of the basin. A comparison between years of strong emission of Rossby waves [START_REF] Rao | Interannual subsurface variability in the tropical Indian Ocean with a special emphasis on the Indian Ocean dipole[END_REF] and the northern Hovmoller diagram from AVISO emphasizes the fact that Kelvin waves end south of 18 N.

The Hovmoller diagram at 16 N shows this alternation of positive and negative anomalies twice a year, whereas at 20 N other wave trains are observed. Using the [START_REF] Fu | Low-frequency variability of the North Pacific Ocean: The roles of boundary-and wind-driven baroclinic Rossby waves[END_REF] method, it is possible to estimate the ability of both the wind stress and coastal Kelvin waves to set Rossby waves. Figure 12 presents the correlation coefficients between the MADT evolution from AVISO with the Rossby waves created by the wind stress field (top) and the coastal waves (bottom). These maps indicate correlation coefficients above 0.6 with the wind forcing from 20 to 22 N and from 23 to 25 N; nevertheless, below 20 N, the winds at the eastern coast have a higher capacity to set up Rossby waves.

Ocean

Interactions of the mesoscale structures

Three different mesoscale features were presented in the previous section: the coastal current, the eddies and the Rossby waves. Hereafter, interactions of these structures with each other and with the atmospheric fluxes are studied.

Rossby wave interactions

A noticeable feature is the intensification of the surface signature in the Hovmoller diagram at 20 N around 64-68 E (see Fig. 11,right). Furthermore, the maps of correlation show low coefficients in the western part of the basin for the Rossby waves forced by coastal waves, west of the topographic ridge and in a region where the mesoscale eddies are more intense. Hereafter, interactions of Rossby waves with the topography and with a dipole of mesoscale eddies are investigated.

www.ocean-

Rossby waves over a seamount

These waves along 21.5 N interact with a topographic ridge between 62 and 64 E. Along 24 N as well, anomalies intensify west of 64 E. To study the influence of topography on Rossby waves, a numerical model of a two-dimensional flow in a bi-periodic domain is used; Rossby waves are initialized with a wavelength of 180 km, propagating over a seamount Chapter5. Mesoscale variability in the Arabian Sea from HYCOM model results and observations : impact on the Persian Gulf Water path with a 30 km radius. Figure 13 shows the flow evolution of the relative vorticity. The Rossby waves deform over the seamount and create a meridional dipole propagating westward; the analytical solution for the flow over the seamount is calculated in the Appendix. It indeed shows the creation of a dipolar perturbation. These results correspond to the propagation observed in the Hovmoller diagram, with a weak anomaly of large wavelengths on the eastern side of the topographic ridge, followed by stronger and shorter anomalies on the western side. The shorter signal is related to the more chaotic nature of the vorticity field resulting from the waveseamount interaction. The stronger amplitude to the west of the ridge, in the Hovmoller diagram (Fig. 11), is due to the vortex-vortex and vortex-wind stress interaction.

Interaction between Rossby waves and a dipole

With the same numerical model, the evolution of a Rossby wave, interacting with a dipole, is studied here in order to understand whether the wave train, propagating westward and encountering the eddies, can reinforce them or not. Different configurations were studied, with different vorticities of the dipole and of the Rossby wave. In the case where the Rossby wave is the strongest, the dipole is either destroyed or deflected northward. In Fig. 14, the Rossby waves deform to recirculate and form eddies, and then merge with the original dipole. Therefore, the earlier interpretation of direct reinforcement of an anticyclone by a Rossby wave (see L 'Hegaret et al., 2013) is not supported by the present simulations. It is in fact a two-step process where the original wave has to deform before it can reinforce the original dipole. Also, the interaction of the dipole with the Rossby wave generates a circular envelope of potential vorticity around it. The sign of the envelope changes at the northern and southern edges of the dipole, with consequent squeezing of the dipole and enhancement of its translational velocity (also observed by [START_REF] Couder | Experimental and numerical study of vortex couples in two-dimensional flows[END_REF].

Eddy merger along the Omani Coastal Current

The northwestern Arabian Sea and the Sea of Oman are filled with mesoscale eddies, originating from the instability of the coastal current, or from the unstable evolution of Rossby waves. These eddies tend to propagate westward, mostly as trains of structures, until they reach the Omani coast. An exception to this westward propagation of eddies is the existence of quasi-stationary eddies near the mouth of the Sea of Oman (25 N) or near Ra's Madrakah (19 N). These eddies can interact in different ways. Indeed, two eddies with opposite-signed vorticity can pair and form a (propagating) dipole. Two eddies with like-signed vorticity can merge if they are close enough.

Figure 15 shows an occurrence of eddy merger near the coast of Oman in summer. The MADT anomaly shows that this merger occurs in the presence of a coastal current, which induces a deformation field on these eddies. In the absence of an external deformation field, eddy merger is known to produce final structures larger than the initial eddies. But in the presence of external shear and strain, filaments are torn away from the eddies and the final structure has a radius similar to that of the initial eddies (about 100 km).

Deformation is maximal at the eddy periphery and between them, while vorticity is maximal in the eddy cores; this latter region is about 50 km in diameter and is characterized by a strong local recirculation. In the eddy cores, the Ertel potential vorticity anomaly is negative, while it is positive in the coastal current. Also, a strip (or a filament) of positive Ertel PVE is advected between the two initial eddies and around the final (merged) eddy. This final eddy has thus an R-structure as described by [START_REF] Morel | Evolution of isolated interior vortices in the ocean[END_REF], whereby the vortex is isolated by a shield of opposite signed, mostly relative, vorticity in its periphery.

The strong ejection of fluid from the merging eddies, into smaller structures, is associated with kinetic energy dissipation and a weaker efficiency of the merging process.

Eddy-surface fluxes (net heat and wind) interaction

Around the year, the surface fluxes have a strong influence on the basin circulation. The wind has an influence on the sea level anomaly through the wind work P = ! ⌧ . ! U g with U g the geostrophic speed at the surface. [START_REF] Roquet | On the patterns of wind power input to the ocean circulation[END_REF] decomposed the wind work into a downward flux of a pressure work term and a lateral divergence term. Only the first term P down can locally modify or create sea surface anomalies. Following Vic et al. (2014), this quantity is integrated here over a mesoscale eddy,

P down = ⇢ 0 g ⌧ ⌘ ! r ^✓ ! ⌧ ⇢ 0 f ◆ , (2)
with h.i = 1 S R R S .
dS, and S the surface of the eddy where the MADT anomaly ⌘ is larger than 5 cm (this value was chosen to have a closed, delimited contour around the structure). [START_REF] Haney | Offshore propagation of eddy kinetic energy in the California Current[END_REF] studied the transfer of kinetic energy from the vertical shear flow to the vertical mean flow in California Current eddies, that is, the baroclinic to barotropic energy conversion (see also Vic et al., 2014). The velocity is decomposed as U = ū + u 0 with ū the barotropic flow and u 0 the baroclinic. The (.) is the vertical mean defined as

(.) = 1 H R 0 H .
dz where H is the depth to the surface where more than 50 % of the kinetic energy is contained. The kinetic energy conversion K 0 K is the sum of two terms

K 0 K ⇠ = ⇢ 0 H ⌧ r ^ ! u 0 ( ūv 0 vu 0 ) (3)
and The first term is related to the relative vorticity of the shear flow and the second to its divergence. In quasi-geostrophic flows the divergence term K 0 K r is expected to be smaller than K 0 K ⇠ (see [START_REF] Haney | Offshore propagation of eddy kinetic energy in the California Current[END_REF]. Five anticyclonic mesoscale eddies were followed in the HYCOM simulation, which stay at the same position and do not interact with other eddies. They had K 0 K ⇠ varying between about 2 ⇥ 10 4 and 4 ⇥ 10 3 W m 2 , and K 0 K r was about 7 ⇥ 10 5 to 5⇥10 4 W m 2 ; thus, there was a ratio of 2 to 8 between the transfer to the barotropic energy due to vorticity and the one due to divergence.

K 0 K r = ⇢ 0 H ⌧ r. ! u 0 ( ūu 0 + vv 0 ) . (4
Figure 16 presents the evolution of the wind work, the transfer term from baroclinic to barotropic kinetic energy and the net heat flux compared to the tendency of the eddies' depth that followed during the spring intermonsoons.

For negative depth tendency (deepening of the eddies) the sum of wind work and K 0 K ⇠ is positive in tendency, mainly due to P down . Other factors such as an eddy merger could occur, but the effect of the wind on the vertical signature of the eddy plays a noticeable role in the barotropization of the eddies. Also, during a deepening, the tendency of the net heat flux is positive, thus indicating a gain of temperature.

When the depth of the eddy remains constant @h @t ⇡ 0 , the term of barotropic energy transfer remains low.

For positive values of the depth tendency (shallowing of the eddies), both positive and negative values of the sum of barotropization terms are found. The strongest values are positive and principally due to the K 0 K ⇠ term. The net heat flux is negative in tendency for shallowing of the eddies (not shown), thus indicating a decreasing temperature.

Evolution of the Persian Gulf Water in the Sea of Oman

The Persian Gulf outflow enters a region of intense mesoscale dynamical activity. The outflow was earlier as- sumed to flow alongshore for most of the year; it was described as a coastal vein during the fall intermonsoon (see Pous et al., 2004b). This is consistent with the presence of the monsoonal coastal currents, but as observed in L'Hegaret et al. ( 2013), fragments of PGW can be found in the Sea of Oman, particularly during the spring intermonsoon under the influence of a dipole off Ra's Al Hamra. Fragments of PGW can also be found in the Arabian Sea (Carton et al., 2012). Using the HYCOM simulation, it is possible to study the evolution of the characteristics, pathway and dilution of the PGW. Four different mechanisms of ejection of the PGW observed in the simulation are presented hereafter. The thickness and width are estimated using the method of Pous et al. (2004b), the limit being the isohaline

⇣ S PGW +S env 2 ⌘ , with
S PGW the salinity of the PGW core and S env the salinity of the surrounding water mass. Also, transport of salt is calculated along the pathway of the outflow. This circulation corresponds to an Eulerian transport, across sections in the Sea of Oman at a specific period, during and after an ejection phenomenon. For each mechanism, the evolution of the MADT anomaly, the deformation field induced by the velocities, and the maximum of salinity at the PGW depth are presented.

Coastal pulse and ejection off Ra's Al Hadd

The focus here is on the evolution of the outflow between Ra's Al Hamra (23.5 N, 58.5 E) and Ra's Al Hadd (22.5 N, 59.8 E) (the PGW evolution in the western Sea of Oman will be discussed in a following subsection). For each year of the simulation, from the beginning of the summer monsoon until December, the outflow is observed to follow the coast of Oman. Nevertheless, short pulses of salt, with salinity anomalies larger than 0.5 psu, are observed two or three times each season for about 15 days (as seen in Fig. 17, left). The MADT anomaly reveals that these pulses are linked to the intensification of a cyclone near Ra's Al Hamra, thus increasing the outflow deformation in this region.

The pulse proceeds as follows: exiting from the strait, PGW is ejected near the Iranian coast, and wraps around an anticyclone located near the Strait of Hormuz; then PGW heads southwards towards the Omani coast as a jet accom- panied by a dipole. The western anticyclone of this dipole remains trapped under the surface anticyclone while the eastern PGW cyclone is elongated along the Omani coast under the influence of the surface cyclone there. This leads to a coastal pulse along the Omani coast between Ra's Al Hamra and Ra's Al Hadd and to its ejection at Ra's Al Hadd.

Figure 18 presents the characteristics of the PGW during and after the pulse. In the left-hand panel, the pulse is associated with an abrupt change in the depth of the outflow and thickness of more than 50 m near kilometers 700 to 900. The salinity above 37.2 psu decreases southeastward at the front of the pulse to the surrounding water salinity, below 36.5 psu. The salt transport is over 2.5 ⇥ 10 8 kg s 1 in the eastern part of the pulse, associated with the strongest velocities. West of the pulse, between kilometers 600 and 700, the outflow recirculates cyclonically, being shallower (near 150 m) and with a thickness of 130 m. The salinity of this cyclone is the same as the pulse and has a salt transport of 1.3 ⇥ 10 8 kg s 1 .

After the pulse, the outflow remains at a depth of 190 m, whereas its thickness decreases from 140 to 110 m. Furthermore, the flow follows the coast, keeping a constant and thin width of 30 km. The salinity also decreases along the coast by 0.1 psu over 200 km. The salt transport is above 10 8 kg s 1 , up to 3 ⇥ 10 8 kg s 1 , strongly depending on the velocities. It then forms a narrow and intense coastal jet.

Offshore from Ra's Al Hadd, the PGW is ejected in summer under the influence of a dipole, the Ra's Al Hadd jet, clearly marked on the deformation map (see Fig. 18). The outflow then wraps around the mesoscale eddies, supplying them with salt, bringing a salinity above 36.9 psu in their core. The filaments at their periphery disappear within a month, being subject to strong mixing. On the map from 21 August, the filament near Ra's Al Hadd recirculates to form a submesoscale structure. This filament formation is observed from the summer monsoon to the fall intermonsoon at a time when the Ra's Al Hadd jet disappears.

Formation of a lee eddy

In winter, the PGW outflow is known to veer southwestward around Ra's Al Hadd. In the HYCOM model, 75 % of the simulated years presented the formation of an anticyclonic lee eddy downstream of Ra's Al Hamra, during the early winter monsoon (for lee eddies, see [START_REF] Verron | Quasigeostrophic flow past a cape in a homogeneous fluid[END_REF]. This formation is related to a strong deformation field around Ra's Al Hamra (see Fig. 19); the outflow is squeezed toward the coast, recirculating in an anticyclonic eddy with a positive MADT anomaly.

Figure 20 (left) shows the characteristics of the PGW during the formation of the lee eddy between 450 and 650 km. The flow becomes wider, about 75 km thicker, to about 180 m, and shallower, with a maximal salinity at 180 m depth. The salinity of the structure does not change from the surrounding water, about 37.2 psu, but the outflow then corresponds to a larger salt transport (about 1.5 ⇥ 10 8 kg s 1 in the interior of the forming lee eddy). East of it, between the 700 and 800 km marks, a filament of PGW is observed, having a large transport, a depth around 180 m and a thickness decreasing by 75 m over 100 km, corresponding to a dilution of 0.2 psu in 100 km.

In the right-hand panel of Fig. 20, the characteristics of the ejected lee eddy and of the filament linking it to the coast are observable downstream of kilometer 750, and the lee eddy is found after the 1000 km position, 3 weeks after its formation. Between the coast and the eddy, the PGW structure is highly variable, beginning with a shallower depth (from 220 to 170 m) and smaller thickness (from 140 to 120 m), as a consequence of a strong and variable deformation field; the outflow is fragmented in filaments. Nevertheless, it keeps a constant salinity at 36.9 psu and a high transport of salt up to 2 ⇥ 10 8 kg s 1 . The lee eddy has stronger anomaly, a salinity around 37 psu, and a salt transport above 3 ⇥ 10 8 kg s 1 . It also has a size comparable to the eddy formed via the deformation of the coastal current with a width of 150 km, a thickness larger than 200 m, and a depth of 180 m.

This structure remains coherent for several months, from December through March, drifting towards the Indian Ocean, slowly eroding, losing about 0.1 psu per month. This is the only feature observed in the model capable of advecting high salinity PGW in the Arabian Sea without being quickly diluted.

Ejection of PGW via a dipole

In spring, the monsoonal winds have a weak intensity, and no coastal surface current is observed in the Sea of Oman. As observed in L' Hegaret et al. (2013), a mesoscale eddy dipole can be found in the basin, with an intense anticyclone east of Ra's Al Hamra and a cyclonic feature west. This dipole is observed in 50 % of the HYCOM simulation results. In the other years, the anticyclone is observed to the west. In Fig. 21, the anticyclone is inducing deformation along the coast (between Ra's Al Hamra and Ra's Al Hadd) and at the center of the basin, advecting the PGW in an anticyclonic motion.

From kilometer 600 on, in Fig. 22 (left), the ejection induces variations of the depth and width of PGW near the coast. The PGW outflow rises and widens before ejection and then equilibrates around 200 m depth and 50 km width, with a salinity of 37.2 psu. Also, the salt transport increases along the coast, up to 2 ⇥ 10 8 kg s 1 , but is weaker, around 0.5 ⇥ 10 8 kg s 1 , in the ejection. The front of the filaments, between kilometers 900 and 1100, presents a difference of 0.6 psu in 100 km and an abrupt change in depth and thickness.

Two weeks later (see Fig. 22, right), from kilometers 400 to 700, a second PGW ejection at Ra's Al Hamra occurs: its transport of PGW is above 2.5⇥10 8 kg s 1 . Nevertheless, the front of this ejection does not reach the coast, as it is occupied by the remaining filament of the previous ejection. Inside the Sea of Oman, after the 700 km mark, the PGW outflow keeps a steady depth of 210 m, a 50 km width and a thickness quickly dropping from 170 to 60 m. The salt transport is low, with 0.5 ⇥ 10 8 kg s 1 , and the PGW vein loses salinity at a rate of about 0.1 psu per week and every 100 km.

Several ejections can occur each spring, similar to the mushroom-like pattern that can be observed on the map of 12 June (Fig. 21). This pattern is associated with recirculations of PGW, part of it being advected eastward, out of the Sea of Oman, whereas another branch veers westward. It is then possible to have, at the same location, two PGW masses with different characteristics, a newer, denser one presenting high salinity, and an older, lighter one that has been subject to dilution and thus presents a lower salinity.

Ejection of PGW along the northern coast near the Strait of Hormuz

In the western Sea of Oman, an ejection of PGW is also observed under the influence of an anticyclonic eddy (see Fig. 23). Directly off Hormuz, the outflow, instead of following the southern coast, is advected northward. The anticyclone induces high deformation of the flow.

Between kilometers 150 and 300 (Fig. 24), the outflow follows the coast, with a depth of 240 m and a salinity of 38 psu

N 150km → 300km → 450km → 600km → 750km → 900km →
PGW pathway after the ejection of a lee eddy associated with the core and edges salinity when reaching the northern coast. Then, the transport of the salt maximal value, near 1.3⇥10 8 kg s 1 , is found at kilometer 250, associated with a deepening and a thickening of the outflow, probably due to an anticyclonic structure observed on the MADT map. After 400 km, this transport becomes negative, due to a convergence of the northern waters. The outflow is ejected from the coast, shallowing to 190 m and with a salinity above 37.4 psu. As for the ejection by a dipole, the front of the ejection presents an important change in salt, near kilometer 550.

The southward ejection reaches the Omani coast with a mushroom-like form, yielding one branch exiting the Sea of Oman, and another one recirculating in the Sea of Oman. A branch of the outflow then wraps around a cyclonic eddy. In the right-hand panel of Fig. 24, from kilometers 420 to 600, the salinity remains constant inside this eddy, around 36. This ejection mechanism is observed primarily during the winter monsoon, with the coastal current advecting the outflowing water northward, but it is also observed in spring and summer, in the presence of an anticyclonic eddy. Finally, this mechanism is not observed in the HYCOM simulation during the fall intermonsoon, where the circulation in the Sea of Oman is dominated by an intense cyclonic eddy.

Salt transport from Hormuz to the Sea of Oman

In the previous panels, for each ejection mechanism, important changes occur in PGW in the first 100-150 km downstream of Hormuz. Exiting the Persian Gulf, at Hormuz, the PGW has salinity between 39.5 and 40, and flows as a bottom current, at 100 m depth and with a width of about 50 km, with an associated salt transport of about 10 8 kg s 1 . At shelf break the flow cascades down on the continental slope and turbulent mixing occurs with the surrounding water, either Downstream in the Sea of Oman, PGW loses 0.1 psu every 100 km; ejection via dipoles or pulses contributes to temporarily halt the dilution, but PGW is then diluted at the periphery of the eddy. When PGW reaches the core of the eddies, the salinity in these cores increases by 0.3 psu, and up to 0.8 psu for the lee eddy. In this study the mesoscale activity in the Arabian Sea and Sea of Oman was assessed by using HYCOM model outputs, measurements from satellites, floats and climatology. Notable similarities were found in the water mass thermohaline characteristics and altimetric signatures between model and data. The surface circulation is highly seasonal under the influence of the wind stress, creating strips of MADT anomaly along the coasts, positive in winter, negative in summer, with branches along the Arabian Peninsula forming the Omani Coastal Current. This surface current deforms to form mesoscale eddies propagating westward with a vertical influence extending below 1000 m depth. Along the coast of India, Rossby waves are radiated under the influence of coastal Kelvin waves, or of wind stress. These Rossby waves have been observed, and in simulation have been found to destabilize and to transform into eddies under the influence of the topography. They can then interact with mesoscale dipoles near the coast of Oman.

The mesoscale eddies interact in different ways across the basin. Two close eddies with the same polarity can merge to form a larger structure; nevertheless, near the coastal current, the final eddy often has a size similar to the earlier ones. Two eddies with opposite polarity advect each other as dipoles. Strong wind favors the deepening and barotropization of the eddies. The net heat fluxes also showed a gain or loss of heat with the deepening or shallowing of the eddies' depth.

The Persian Gulf outflow is impacted by the mesoscale eddies, influencing its pathway along the Sea of Oman; PGW ejection from the coast occurs, from the Ra's Al Hadd cape in summer, from Ra's Al Hamra in spring and fall, and along the northern coast in winter. The thermohaline characteristics are also changed by recirculation in the western part of the basin, by the generation of eddies or by the breaking of the core, these phenomena leading to thermohaline mixing. The injection of PGW inside eddies allows this water mass to keep a stronger thermohaline signature across the Arabian Sea, filaments being more subject to erosion, whereas other structures, such as lee eddies in winter, keep higher salinity for several months. Note that the salinity of PGW decreases mostly in the first 150 km from the Strait of Hormuz as a consequence of cascading or of entrainment on the bottom slope.

Thus, this study provides a description of the path of PGW outflow from its exit from the Strait of Hormuz to its dispersion in the Arabian Sea. Several types of mesoscale eddies containing PGW (lee eddies, isolated eddies, dipoles) were found, as well as filaments. Further investigation of these submesoscale structures using in situ data of the Phys-Indien experiments, and the analysis of fine-scale intrusions at thermohaline fronts, will be the subject of a forthcoming paper.

Ocean

Further studies : Seasonal variability

To complement this study, we describe here the seasonal varialibity for the Arabian Sea and the Sea of Oman, using the decomposition in EOFs from the altimetric maps.

Summer monsoon

During the summer monsoon, the interannually averaged MADT anomaly indicates a strip of low sea surface near the coast, due to the o↵shore Ekman transport. The first EOF (see figure 5.1, left) exhibit this coastal anomaly as 29% of the variability with a recurrence every year. This signature is also present on the second EOF (10%) along the southwestern coast. Furthermore, on the three first modes of variability, several eddies are found along the coast of Oman. In the Sea of Oman, two opposite-signed MADT anomalies are established in summer : a first one, positive, lies between Hormuz and Ra's Al Hamra, West of the basin, as observed on the two first EOF. A second altimetric anomaly, negative, lies between Ra's Al Hamra and Ra's Al Hadd, forming a stationary cyclonic eddy at the end of the summer monsoon. This negative anomaly is observed on the three first EOF, but on the first it is melt with the coastal "belt".

In summary, the main features of the summer monsoon, from a surface signature point of view, is the coastal belt, that serves as a guide for the Oman Coastal Current; the dipole North and South of Ra's Al Hadd that creates the monsoonal "Ra's Al Hadd jet"; and the western mesoscale eddies that advect the water masses in and out of the di↵erents basins.

Fall inter monsoon

During the fall inter monsoon, the anticyclonic eddy near Ra's Al Hadd loses intensity and is advected southwestward towards the coast. The cyclonic eddies are advected southwestward, except the one at the mouth of the Sea of Oman that reintensifies. These tendencies are underlines on figure 5.1 (right), on the first EOF (21%).

The second and third EOFs (8.6% and 8.2%) point out the presence of the strong eddies extending far more east in the basin than during the summer, nvertheless keeping stronger intensities on the western side. With the exception of the cyclonic eddy in the mouth of the Sea of Oman, the others eddies have an alternating rotation direction depending of the year.

Along the eastern coast, the onset of the altimetric "belt" appears on the first EOF.

Winter monsoon

During the Winter monsoon, winds blow southwestward along the Omani coast and southward along the Indian coast. It creates a downwelling and a positive MADT anomaly along the coast of Oman. The crest of MADT along the coast of India is likely to be related to a Kelvin wave. In November, a positive altimetric anomaly forced by the winds hugs 5.11. Further studies : Seasonal variability the coasts, replacing the negative anomaly of summer and autumn, and several cyclonic eddies lie o↵ the Omani coast. In November and December, this positive anomaly intensifies, destabilizes, forms meanders and advects the cyclonic eddies southwestward. In particular, the cyclonic eddy at the mouth of the Sea of Oman is advected southward of Ra's Al Hadd and then remains stationary all season long; a second cyclonic eddy exists o↵ Ra's Madrakah.

The first EOF mode (34%) shows clearly the strong variability on the west coast but also the positive anomaly along the coasts. The second and third mode, respectively (6.7% and 5.2%) point out that the maximum of variability lies at the mouth of the Sea of Oman and along the coast of Oman. The eddy North of Ra's Al Hadd is cyclonic, but its intensity varies every year. 1997 and 2009 being years with the lowest signature. This monsoon season is strongly alike the summer, with coastal current, but this time southward along the coast of Oman. This "belt" and the cyclone at the exit of the Sea of Oman constrain the Persian Gulf Water and the other water masses along the southern coast. Out, in the Arabian Sea, the surface circulation is dominated by mesoscale gyres.

Spring inter monsoon

During the spring inter monsoon, the positive MADT anomalies along the coasts (the "belt") forms meanders and anticyclonic eddies, first along the north coast, then along the Omani coast. These anticyclones and the cyclone o↵ Ra's Madrakah propagates southwestward. The cyclonic eddy o↵ Ra's Al Hadd remains stationary, coupled to an anticyclone north of it, in the Sea of Oman. Furthermore, an anticyclone is generated in this seasons and remains until the summer monsoon, northeast of Ra's Madrakah, at 59 o E-

60 o E 19 o N-20 o N.
The two first EOF modes (see figure 5.2, right), with representativities of 24% and 11%, show the strong variability along the southern coast of Oman, and north of Ra's Al Hadd. Their signature is always the same for the years recorded, emphasizing the recurrence of these structures. The third one also point out structures in the Sea of Oman, particularly a cyclone o↵shore Ra's Al Hamra, and along the couthern coast of Oman.

In the next section we focus on a specific season, the Spring 2011, du ring the Phys-Indien experiment, in order to study, with a higher resolution these sstrucutures and their impact on the water masses, from the mesoscale to the sub mesoscale.

Motivations for the research

In the previous chapter, the mesoscale circulation in the northern Arabian Sea and Sea of Oman was studied, the evolution of mesoscale eddies, as well as the surface and deep circulation induced by these eddies or by other seasonal features. Also, an emphasis was put on the thermohaline characteristics, seasonal spreading and eddy interactions of the Persian Gulf Water outflow; this allowed us to have a description of its evolution over an average year.

Here, the focus is put an even finer description of the interactions between the mesoscale structures and the Persian Gulf outflow, using in situ data from the Phys-Indien experiment during the spring 2011 inter monsoon. A description of the data and processing is Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring given in chapter 2.

During this experiment, the layer of Persian Gulf Water (PGW) was observed in the Arabian Sea and in the Sea of Oman, with varying characteristics inside each basin; submesoscale structures, filaments and lenses were also sampled and recorded.

This study describes first the regional situation from March 2011, using the same tools as in the previous sections, in order to fit the period of measurements with the annual cycle and to compare the spring 2011 with the springs of other years. From this, a few features are expected to appear in the data, such as the ejection of PGW near Ra's Al Hamra, as described in chapter 4. Then an analysis of the mesoscale structures is given for both basins, with their impact on the water masses and the circulation in March 2011. From this, an emphasis is given on the Persian Gulf Water characteristics, showing strong variations depending on its location. Finally, a a submesoscale lens containing highly saline Persian Gulf Water, observed o↵ Ra's Al Hadd, is studied, from its thermohaline and velocity structures, to its formation and evolution.

This study is submitted to the Journal of Physical Oceanography, hereafter the submitted version is presented.

Introduction

The Indian Ocean, the third tropical basin in size, is bounded north by the Asian landmass.

The orography of this landmass leads to the existence of monsoons, which strongly influence the regional oceanic circulation. The northwestern part of the Indian Ocean is comprised of different sub-basins, each with specific geographic and climatic characteristics. This study focuses on two of them, the Sea of Oman and the Arabian Sea (see figure 1). The Sea of Oman (or Gulf of Oman) connects the Persian Gulf to the Arabian Sea; it deepens along its zonal axis, to 3400 m depth in its eastern part; it also widens from the Strait of Hormuz to its mouth, at Ra's Al Hadd. The northwestern Arabian Sea has a floor below 3500 m depth and is crossed from north-east to southwest by the Owen Fracture Zone, with diving and rising of the topography along the fault. Along the Omani coast a plateau of 200 m depth extends from Ra's Al Hadd to Ra's Madrakah (see again figure 1 for locations).

The surface circulation around the Arabian Peninsula is dominated by mesoscale eddies formed via different processes, depending on the location and on the time of the year (see Fischer et al. (2002)). Al Saafani et al. (2007) identified that such eddies in the Gulf of Aden, which have a westward propagation, are generated by Rossby waves emitted from the Indian coast or in the interior of the basin. This Rossby wave mechanism is also present in the northern Arabian Sea, these waves being forced by wind and by coastal Kelvin waves (see L 'Hégaret et al. (2015)).

Another formation process of mesoscale eddies is the instability of alongshore currents, during the monsoon seasons (see again L'Hégaret et al. ( 2015)). In summer (winter), these currents are intense and have a negative (positive) altimetric signature as coastal "belts", which grow under the influence of the wind stress curl. These coastal currents then form meanders and eddies, which have a radius comparable with, or slightly larger than the first baroclinic radius of deformation 3 Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring (about 40 km in the region, see [START_REF] Chelton | Geographical variability of the first baroclinic rossby radius of deformation[END_REF]). These eddies have a strong velocity, above 60 cm.s 1 , near the surface, and a vertical influence on the water masses at depth (see [START_REF] Bower | Mesoscale eddies in the gulf of aden and their impact on the spreading of red sea outflow water[END_REF] and Carton et al. (2012)).

Over the Persian Gulf, intense evaporation (see Privett (1959) and Meshal and Hassan (1986)), steady winds and little fresh water inflow (through precipitation and river, see Reynolds (1993)), lead to the formation of highly saline water in the northwestern and southwestern parts of this semi-enclosed sea. This saline water then flows into the Sea of Oman via the Strait of Hormuz. In the Sea of Oman, this dense outflow of Persian Gulf Water (hereafter PGW), equilibrates around 250 meters depth, mixing with surrounding and fresher Indian Ocean Surface Water. Near the Strait, PGW has high thermohaline characteristics, with salinity above 39 psu and temperature above 22 o C.

In the Sea of Oman, the upper 100 meters of the water column are occupied by the Arabian Sea High Salinity Water, a water mass formed in the northeastern Arabian Sea in winter (see Kumar and Prasad (1999)), which spreads in the western basin through the year. Its thermohaline characteristics in the Sea of Oman and northwestern Arabian Sea are salinity around 36.6 psu and temperature above 22 o C.

In the past, few dedicated cruises provided observations to describe the PGW pathway out of the Persian Gulf and its variations in the Sea of Oman. This PGW outflow is usually presented as a southeastward flow, along the coast of Oman (see [START_REF] Premchand | Watermass structure in the western indian ocean, part ii, the spreading and transformation of the persian gulf water[END_REF]). In October-November (fall intermonsoon), the GOGP99 experiment sampled the Persian Gulf outflow and identified it as a coastal layer, extending to the south of the Sea of Oman (see Pous et al. (2004)).

More recently, ARGO floats (see Carton et al. (2012) This paper main objective is twofold. First, to describe the mesoscale structures in spring 2011 and their induced circulation on the water masses. Second, to concentrate on the submesoscale fragments detached from or by the mesoscale eddies, and then, on the nature, structure, recurrence and possible role of such fragments.

Data & Method

a. Climatologies and long time observations

Altimetric maps are obtained from the AVISO satellite data center; these maps are created via data merger from Jason 1, Jason 2 and Envisat measurements (see Ducet et al. (2000)). The alongtrack data are then interpolated on a 1/4 ⇥ 1/4 degree Mercator grid with a daily value. The error on along-track measurements of sea level are 3-4 cm (see Fu and Cazenave (2000)). Adding a 22-year mean of the sea surface topography, mean absolute dynamic topography (MADT) is generated. In each MADT map, an spatial (and instantaneous) average is subtracted, to obtain anomalies of the sea surface elevation (MADT anomalies). The wind stress and wind stress curl are obtained from the ASCAT database of Ifremer, with a daily mean and a 1/4 ⇥ 1/4 degree horizontal resolution, from 2007 to 2014.

To acccurately describe the water masses, the Generalized Digital Environment Model (GDEM) climatology is used (http://www.usgodae.org/). It provides monthly mean salinity and temperature maps, with 72 vertical levels and a 1/4 ⇥ 1/4 degree horizontal resolution. These maps are modified by merging measurements from ARGO floats (from the ANDRO database see Ollitrault and Rannou (2013)). About 300 floats, with periods varying between 5 and 10 days, sampled the Arabian Sea and between 2002 and 2014. This provides enough profiles for monthly corrections of the climatology in each sub-basin. This correction is calculated as follows: first, each float measurement is converted into an anomaly by substracting the climatological value at the given time and location. Second, this pointwise anomaly was extended spatially with an isotropic Gaussian correlation function, weighted by the local number of data (the radius of the Gaussian was chosen as L = 50 km, typical of a mesoscale eddy). This correlation function was normalized to unity, where data lay. Third, the resulting anomaly is added to the climatology (see figure 2, left and center). The third panel of this figure (right) indicates the number of profiles, with their radii of correlation.

b. The Phys-Indien 2011 measurements

The Phys-Indien 2011 experiment measured the circulation and water masses in the sub-basins around the Arabian Peninsula, from the Red Sea to the Persian Gulf from late February to April 2011, using many devices (CTD on Seasoar, CTD and L-ADCP stations, XBT, XCTD casts, en-route VM-ADCP, release of Surdrift buoys and of Provor floats, collection of meteorological data; see the sections on figure 3). This study focuses on the measurements in the northwestern 6 6.3. Data and methods Arabian Sea and in the Sea of Oman, in March 2011, with a high-resolution snapshot of the thermohaline and dynamical characteristics of the region obtained by the experiment. In particular, these sections crossed mesoscale eddies and the PGW outflow and fragments.

The thermohaline sections were obtained with two CTD's on a SeaSoar, a towed device.

Pressure, temperature and conductivity are thus obtained in the upper 350 meters of the water column with accuracies of 10 3 o C, 3 10 4 S.m 1 and 0.015% of the pressure value). Salinity is calculated from temperature and conductivity. In this region, strong horizontal and vertical thermohaline gradients occur, which lead to biases in temperature and salinity data on seasoars.

These biases are due to thermal inertia of the sensors, most occur between the lowering and rising of the device, which result in a delay between the conductivity and temperature measurements (see Barth et al. (1996)). These errors can be corrected by applying a couple of coefficients, a first one correcting the amplitude of the measurements between the rising and lowering of the SeaSoar, and a second coefficient correcting the time delay (see Lueck and Picklo (1990)). Once corrected, the SeaSoar measurements are validated using independent CTD station or XCTD casts data (such stations or casts are achieved along the SeaSoar transects). The relative residual error of the corrected SeaSoar salinity remains below 0.5% of the CTD's.

The horizontal velocity was obtained with a 38 kHz Vessel Mounted -ADCP (VM-ADCP hereafter); this device provides data most often from the surface to about 1000 meter depth, with an accuracy of 5 10 3 m.s 1 (in a few instances, deeper reach, down to 1600m, is obtained for the VM-ADCP). This device produces high resolution measurements but is sensitive to noises due to small scale dynamical processes or to biological activity. A low pass filter is applied to the signal keeping the structures with a size superior to 3 km, to focus on the submesoscale and 7

Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring mesoscale processes. Another VM-ADCP, with 150 kHz frequency, was activated but its data are not shown here (due to too shallow reach). These data are used nevertheless to validate the 38 kHz ADCP measurements in the upper ocean. The VM-ADCP data are filtered, but a few sections were interrupted or were too noisy, thus having blanks. No interpolation is then carried out and the blanks are displayed on the figures. This may also be the case of a few SeaSoar transects.

During Phys-Indien, five drogued Surdrift buoys were deployed and programmed for 180 days of recording. These surface buoys are connected to a large holey-sock drogue by a thin Kevlar cable, 80 to 250 m long. A test on the acceleration of each buoy was applied to determine a possible loss of the drogue. These buoys were positioned by Argos, their trajectories were sampled every hour and a thermistor sensor gave the surface temperature at each recording. Also, 6 PROVOR floats were deployed during this cruise. They were positioned via Argos when they surfaced at the end of each 5-day cycle. These floats were equipped with a CTD probe providing temperature, conductivity and pressure with an accuracy of 0.01 o C, 10 3 S.m 1 and 1 dbar respectively. Their parking depth was programmed to 700 dbar instead of 1000 dbar; they dived to 2000 dbar every 5 days and acquired data while rising to the surface, where these data were transmitted.

c. Thermodynamical and dynamical quantities derived from the measurements

To describe the structure of the mesoscale eddies, surface maps of MADT anomaly are computed at the period and location of Phys-Indien measurements; from this anomaly, surface geostrophic velocities (U,V ), relative vorticity and the Okubo Weiss quantity are obtained through derivations. The Okubo-Weiss quantity is defined as a difference ,(in norm) between deformation 8 6.3. Data and methods and relative vorticity

OW = s 2 strain + s 2 shear w 2
with the shear,

s shear = ∂V ∂ x + ∂U ∂ y the strain s strain = ∂U ∂ x ∂V ∂ y
and the relative vorticity

w = ∂V ∂ x ∂U ∂ y .
The Okubo-Weiss quantity is positive in regions where deformation dominates rotation, and it is negative where vorticity dominates. It is calculated in the upper 350 m of the water column using density from the ARGO-GDEM climatology.

From the thermohaline data, density and spiciness are calculated along the sections. In-situ density is obtained by a state equation from IOC and IAPSO (2010). Density anomaly s 0 is displayed. Using T 0 , S 0 and r 0 , the reference temperature, salinity and density, of 20 o C, 37 psu and 998 kg.m 3 , spiciness g is calculated :

g = g 0 [1 + a (T T 0 ) + b (S S 0 )].
Using these SeaSoar density fields and surface velocity calculated from MADT from AVISO, 9

Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring sections of geostrophic velocities are computed in order to compare them with the VM-ADCP velocity measurements.

Using velocities from the VM-ADCP measurements and density from the SeaSoar it is possible to calculate a two-dimensional Ertel potential vorticity (EPV) (see [START_REF] Hoskins | The role of potential vorticity in symmetric stability and instability[END_REF]).

EPV = f ( f + ∂V g ∂ x ) ∂ b ∂ z f ∂V g ∂ z ∂ b ∂ x with b = ( g r 0 )r the buoyancy.
Note that these calculations lead to some noise in the results; the figures presented in the text are slightly filtered for legibility.

Mesoscale situation from late 2010 to mid 2011

The spring inter monsoon extends on average from February to May in the Arabian Sea. During this period the wind stress over the basin is low, between the two local maxima of the winter and summer monsoon. The wind stress curl is positive in winter and negative in summer with strongest values. Intermonsoon periods show a change in sign of the curl and weak winds.

The wind stress from summer 2010 to summer 2011 (see figure 4), followed the climatological In May-June 2011, the sea surface warmed up, cyclone C1 weakens, A2 intensifies in relation with the increasing negative wind stress curl and the alongshore current formed again in response to the onset of the summer monsoon.

The observed evolution of the structures in the region during the spring 2011 follows that of an average spring inter monsoon, with large structures dominating the surface circulation and a strong anticylonic signature in the eastern Sea of Oman. The eddies influence the distribution of the sea surface temperature, intensifying the thermal front near Ra's al Hadd (and later on advecting cold water offshore); they also have an influence at depth, in particular the anticyclone in the eastern Sea of Oman can eject the PGW offcoast towards the northern Iranian coast, and advects these waters around it (already mentioned in Carton et al. (2012)). This will be evidenced now with the Physindien experiment data.

a. Arabian Sea

The Arabian Sea (AS) section studied here is located in the Arabian Sea, south of Ra's Al

Hadd, from 16 o N to 22 o N almost along 60 o E; it was carried out from March 16 th to 19 th 2011.

The MADT anomaly map (figure 6, top, left) and the relative vorticity map (figure 6, top, left) indicate that section AS intersected, from south to north, a cyclone (C2), and a stronger anticyclone-cyclone pair (A2-C1). Cyclone C2 is part of a three cyclone compound, south of A2-C1. These cyclones were intensified in January 2011 by a westward propagating, negative MADT anomaly, generated earlier from the east.

The map of surface relative vorticity confirms that C2 was weak (w < 10 5 s 1 ). Cyclone C1 was the strongest eddy (w > 3 10 5 s 1 ) and the anticyclone A2 had w ⇠ 1.5 10 5 s 1 (figure 6, top, right). The fastest velocities (60 cm.s 1 ) were located on a ring around cyclone C1 and on the jet between A2 and C1 (figure 6, bottom, right). A2 also presented strong surface velocities, above 35 cm.s 1 .

The Okubo-Weiss field (figure 6, bottom, left) indicates maximal deformation around cyclone C1 and on the A2-C1 jet.

The velocity structure of the eddy has a strong geostrophic component (see figure 7, top). The measurements indicated stronger values in the eddies, between 100 m and 150 m depth with 13

Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring maximum velocity, above 80 cm.s 1 , than those computed geostrophically. The front between the cyclone and anticyclone is observed at the same position in both velocity sections. Furthermore this velocity remained noticeable below 600 m depth, with values above 10 cm.s 1 (not shown).

This confirms the deep dynamical influence of these eddies and their barotropic component.

The EPV and Okubo-Weiss help identify the coherent structures (see figure 7, center). The maximal values are found in the core of the eddies, positive for the cyclones C1 and C2 and negative for the anticyclone A2, between 50 and 150 m depth. Furthermore, the position of these maxima are correlated with a negative Okubo-Weiss quantity, meaning that rotation dominates deformation, whereas the periphery of the eddies are characterized by predominance of deformation over rotation.

The thermohaline sections (figure 7 The depths and characteristics of these three water masses are strongly influenced by the mesoscale eddies, as revealed on the density and spiciness sections(figure 7, bottom). The anticyclone and the cyclone induced a vertical stretching or squeezing of the isopycnals. The depth of ASHSW varied thus from 150 m depth, below the anticyclone, to 80 m depth below the cyclone. Similarly, PGW lied between 250 and 350 m depth under anticyclone A2, 50 m deeper than out of the eddies. PGW shoaled under cyclone C1, between 180 and 300 m depth.

On the profiles below the anticyclone (red circles) and below the cyclone (blue circles) on diagram figure 8), PGW has a density s 0 = 26kg.m 3 below the anticyclone, whereas it is denser under the cyclone with s 0 = 26.5 kg.m 3 . Thus, PGW below these eddies may have been captured

at different times and locations, since the PGW outflow from the Persian Gulf presents varying characteristics with time.

The spiciness section (see figure 7, bottom) characterizes the thermohaline gradients across isopycnals, revealing the boundaries between the water masses. The most relevant feature of spiciness on the AS section lies around the mesoscale eddies: layers of cold and fresh water wrap In summary, in the Arabian Sea, the surface circulation is dominated by mesoscale eddies, primarily along the coast. Their strong vorticity in the upper layer separates ASHSW from water masses below. These eddies also have a deep vertical influence so that thus trap horizontally, displace vertically and mix the water masses below them. Under the anticyclone, the temperature of PGW is 2 o C colder under the cyclone. The difference in temperature can either result from a vertically heat flux between the eddies and these deeper water masses, or from the advection of 15

Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring a specific part only (e.g. the upper or lower part), of these water masses, under the eddies. They can also wrap IOCW and PGW around them, as layers. Submesoscale structures (filaments), with homogeneous water masses, are advected between the main eddies.

b. Sea of Oman

From March 22 nd to 30 th the PhysIndien experiment performed cross-sections in the Sea of Oman. A composite section (SO section) crossing this basin zonally, is presented and described here.

The MADT map above SO shows a positive anomaly (see figure 6), formed from a westward propagating signal, which entered the Sea of Oman in December 2010. The relative vorticity map reveals that the western part of the basin is occupied by smaller eddies. Thus, from west to east, SO crossed three alternating anticyclones and cyclones. The most energetic eddy, A1, laid in the eastern part of the Sea of Oman, with the highest velocities, up to 25 cm.s 1 . The Okubo-Weiss quantity is negative in the center of A1, indicating a dominance of rotation, but positive in the western part, undergoing deformation.

As in the Arabian Sea, the maximal velocities of the eddies laid in the upper 150 meters (see figure 9, top), but a secondary maximum was found west of 57. The EPV field between 50 and 100 m depth is negative, as expected from the anticyclonic surface signature in the MADT, a positive maximum is found in the westernmost part of the section, characteristic of cyclonic eddies (see figure 9, middle). At depth, a positive peak in the EPV field is observed between 58 and 58.7 o E, where Okubo-Weiss shows that deformation dominated over rotation. Furthermore, the spacing between the isopycnals, and the velocity signature indicates a anticyclonic motion, underlining the presence a of lens.

The thermohaline section (figure 9, bottom) presents, as in the Arabian Sea, three water masses.

The The spiciness also underlines the turbulent mixing, with injection of colder water and fresher water around the patches of PGW, resulting in a fragmented outflow water mass. Inside this patch, the layer can be vertically discompose, with the most saline water between 220 and 320 m depth, with s 0 = 26.2 kg.m 3 , and a second layer, between 150 and 230 m depth, with s 0 = 25.8 kg.m 3 . Around the eastern anticyclone, with the densest PGW at depth, an injection of saltier and warmer water inside this layer is observed.

The circulation in the Sea of Oman shows strong resemblance with that in the Arabian Sea.

Mesoscale eddies, with lower intensity, dominate the flow at the surface and at depth. The water masses immersions depend of the pressure anomalies of the upper ocean layer. In the Sea of Oman the outflowing PGW is subject to strong mixing, and is fragmented, with detached rotating fragments at depth, and more isolated patches above. To the east of the gulf, beyond the strongest velocity front between surface eddies, the outflow is not clearly observed; the few Persian Gulf The spreading of the PGW during the Phys-Indien experiment is of interest, due to its variability in position, depth and characteristics. Hereafter the focus is put on its recording during the experiment.

PGW characteristics, pathway and submesoscale structures

On the q /S diagrams of figures 8 and 10, an important feature of the PGW is the peaks in salinity at two different density ranges. A first one, lighter and fresher, with s 0 between 25.8 and 26 kg.m 3 is found on the profiles in the western Sea of Oman and below the anticyclones A1 and A3 with a salinity around 37 psu. This peak is also observed with diluted characteristics inside the core of the eddies in the western Sea of Oman and Arabian Sea and with salinity below 36.4

psu. A second peak, with higher salinity and denser, is found in the eastern Sea of Oman with a salinity of 37.9 psu, below the eddies with a salinity varying from 36. 12) are computed to follow the pathways of the PGW outflows. The Eulerian transports can be related to the velocity sections and to the altimetric maps. Indeed, in the Sea of Oman, between 58.5 and 60 o E, the volumic transport was anticyclonic, and also between 57 and 58.5 o E. This corresponds to A1 and A3.

In the Arabian Sea, the strong jet between C1 and A2 led to the 6.43 Sv transport eastward at 20.5 o N, while an opposite westward jet between A2 and C2 at 18 o N transported 2.77 Sv.

In the Sea of Oman, the front already seen in the SO section appeared with warm and salty PGW west of 59 o E, and fresher waters with patches of higher spiciness east of 59 o E. This front was created by the strong anticyclone A1, which stopped the zonal spreading of the PGW and advected it northward. The penetrating IOCW is also blocked from the east by the front, as observed on the recording of an ARGO float (number 2901370) in March 2011 (see figure 13). Differences between the altimetric map and the transport map can be due to the effect of bathymetry or of thermocline density gradients.

In In the Arabian Sea, the PGW motion around the eddies C1 and A2 was similar to that in the Sea of Oman but with stronger dilution, by more than 1 psu (see figure 8). In the A2 anticyclone, the PGW layer was warmer (see figure 11, top, left), with temperature above 19.5 o C,a priori due to heat transfer from the ASHSW layer in this intense eddy. South of 18 o N, the PGW was strongly diluted and mixed with the IOCW.

21

Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring

The velocity section (figure 14, top) shows that the signature of the cyclone extends below 1000 m depth in the southern part of the section and below 400 m depth in the northern part.

Between 22.1 o N and 22.4 o N and 250 and 400 m depth, an anticyclonic motion is observed on the zonal velocity section with speed of about 20 cm.s 1 , but strongly dominated by the velocity of the surface eddy. The Okubo-Weiss field structure (figure 14, center) indicates that the first 400 meters of the ocean were dominated by the strong deformation field between the two eddies and that a maximum of deformation is observed at the depth of the lens. The presence of this lens is also indicated on the EPV section by the anomaly.

The thermohaline section confirms the presence of a lens shaped structure between 250 and 350 m depth, with a diameter of 25 km, a temperature of 19 o C and salinity over 37.3 psu at its center.

On a perpendicular cross section (not presented here), the lens has a diameter of 33 km, indicating that it is elliptical, this is confirmed by a third section (not shown floats recorded temperature and salinity with a surfacing every 5 days. All these floats followed a northwestward trajectory for the first 4 days of measurements. After this period, the three ARGO floats were ejected, as revealed by their recorded salinity; so were the surdrifts which then followed anticylone A3, north of Ra's Al Hadd. One surdrift buoy lost its drogue and performed inertial loops. This inability to track the lens for long underlines the strong deformation that it was subjected to. [START_REF] Ruddick | The production and dissipation of compensated thermohaline variance by mesoscale stirring[END_REF] studied the deformation of a vortex lens by shear and strain.

As the strain increases, the lens become more elliptical and unsteady, before either breaking up or readjusting. The surdrift loops reducing in size and velocity, and then following the mesoscale anticyclonic flow are coherent with a break up of the submesoscale PGW lens. The energetic mesoscale eddies have a strong vertical influence and deepen or shallow the water masses below them (ASHSW, IOCW and PGW). Inside these eddies cores, these water masses keep the thermohaline characteristics at the time of their trapping; PGW filaments wrapping around these eddies are subject to strong dilution. Besides, injections, primarily of cold and fresh IOCW, occur around the eddies, thus inducing mixing, and dilution of the highly saline waters.

In the western Sea of Oman, the PGW outflow appears fragmented, forming small eddies, filaments and a few isolated patches. Two layers of PGW, with different densities, from the winter monsoon and earlier mixed PGW, were observed at the same location, due to the anticyclonic recirculation in the western basin. The PGW outflow was not observed in the measurements along the coast between Ras Al Hamra and Ra's al Hadd, with the exception of a few small coastal patches; PGW was advected north, around anticyclones A1 and A3, slowly diluted along its pathway, with a salinity below 37 psu.

A θ/S profiles in the Arabian Sea

Conclusion and Perspectives

Conclusion

The first aim of this thesis was to study the circulation in the Arabian Sea and in the adjacent basins over the di↵erent seasons. The monsoon dominates the dynamics of the Indian Ocean, thus creating a complex circulation, particularly near the Gulf Of Aden and the Sea of Oman where a study from a mesoscale perspective was of interest.

Using monthly climatology, modified with the ARGO floats circulating in the region, mesoscale eddies appeared to dominate the circulation for all seasons, particularly along in the coastal western Arabian Sea and in its two adjacent basins. However, coastal currents also advects these floats for some periods of time. Comparisons of the altimetry and the trajectories of the floats display high correlation, thus indicating a deep signature of these mesoscale eddies.

The second aim of this thesis was to investigate the impact of the mesoscale circulation on the water mass, particularly the highly saline waters formed in the Red Sea and the Persian Gulf. The ARGO float trajectories showed that these outflows can be advected around but also inside the mesoscale eddies, thus having di↵erent thermohaline characteristics.

An emphasis was put on a seasonal feature that appeared recurrently on the altimetry in the Sea of Oman; a dipole during the spring inter monsoon. The ARGO floats, as in the Arabian Sea presented correlations between their trajectories and this structure. An intensification of the anticyclonic part of the dipole was linked with a change of the wind stress curl during this season. As well, Rossby waves were observed to propagates westward all year long, with a positive anomaly arriving at the position of the anticyclone during this period.

In previous studies, the Persian Gulf outflow had been observed to flow along the southern coast in the Sea of Oman during the monsoon seasons. With the ARGO floats, patches of Persian Gulf Water were observed in the northern part of the basin, around the mesoscale eddies. Furthermore, Persian Gulf Water was also observed, with di↵erent characteristics, inside the eddies.

From this, a study of the mesoscale variability over the northern Arabian Sea and the Sea of Oman was carried out over each season, using observations and numerical results of HYCOM. The monsoon winds induce the formation of a coastal belt of altimetry, with a positive or negative anomaly, in winter or summer. The coastal currents take place along these anomalies, which meander and produce mesoscale eddies, along the coasts. These eddies have sizes comparable with the first baroclinic radius of deformation. Along the western coast, these structures propagate southwestward, some of them remaining steady for several months, particularly near the mouth of the Sea of Oman. On the eastern side, Conclusion & Perspectives these structures propagate westward; as well Rossby waves are emitted from the eastern coast of India but the topographic mounts in the basin destabilize them to form eddies. At depth, these eddies have a strong barotropic signature, reinforced by the wind stress curl.

An emphasis on the Persian Gulf outflow in the Sea of Oman reveals the strong seasonality of its characteristics and its pathway through the basin. From summer through winter, the outflow most often flows along the southern coast but it can wrap around the mesoscale eddies or forms coastal lee eddies. In the eastern part of the Sea, and particularly in spring, recirculations are observed, and the dipole creates an o↵shore ejection near Ra's Al Hamra, advecting water along the northern coast, then recirculating or exiting the Sea of Oman.

Then, using the dataset from the Phys-Indien 2011 experiment, a study of mesoscale eddies and of Persian Gulf Water with higher vertical and horizontal resolutions was performed along the northwestern coast of the Arabian Peninsula in March 2011. The vertical influence of the mesoscale eddies reveals that the Persian Gulf water is captured below them and advected along their life cycle. Nevertheless, lateral injections of colder and fresher oceanic water enhance the dilution of the highly saline waters, primarily in the Arabian Sea.

Still in spring 2011, the Persian Gulf Water in the Sea of Oman recirculates in its eastern part, thus spreading at two di↵erent densities, a lighter one from the winter and a denser one from spring. The Persian Gulf Water fragment appears as highly saline small eddies, filaments and isolated patches, that are ejected o↵shore under the influence of the Ra's Al Hamra dipole.

A submesoscale lens of dense Persian Gulf Water was observed in the Arabian Sea and was related with the coastal ejection of outflow in the Sea of Oman. During the period of measurements, this lens appeared to be elliptic, deformed by the mesoscale eddies neighbouring it, thus inducing its breaking.

In summary, to answer the questions listed in the introduction : "-Which features dominates the mesoscale circulation in the Arabian Sea ?" The observations, through ARGO floats and satellite altimetry, as well as the numerical simulations, indicate the strong presence and importance of mesoscale eddies in the Arabian Sea, with the most intense gyres along the western coast.

"-How do they form and what are their structures ?" They form from the coastal current that appears during each monsoon. This current deforms, with a wavelength comparable to the first baroclinic radius, into the mesoscale gyres. They have a radius of about 75-100 km and a barotropic signature with an influence below 1000 m depth.

"-How does the monsoon seasonality impact them ?" The seasonality is clearly observed along the western coast of Oman, where the most stationary vortex structures have an opposite vorticity sign in summer or in winter. As well the inter monsoon period shows the presence of more mesoscale eddies than during the monsoons. Finally the wind stress, through its curl, impacts the barotropization of these eddies.

"-What is the interaction between these mesoscale structures and the water masses of the region ?" The eddies, with their strong vertical structures, capture water masses inside and below them. They advect and transport these water masses southwestward, along their trajectories. Also, surrounding water are advected around these gyres, increasing their deformations. As well, fresher oceanic water is injected inside the eddies, thus enhancing dilution of the saltier waters.

"-What is the evolution of the Persian Gulf outflow pathway in the Sea of Oman throughout a whole year ?" The Persian Gulf Water is influenced by the gyres. During the monsoon, it flows along the coast of Oman, then being advected around the gyres, ejected at the Ra's Al Hadd cape, or forming coastal lee eddies (particularly during the winter monsoon). In spring, a cyclonical recirculation is observed in the western Sea of Oman, and an strong anticyclone creates a front, the PGW being ejected o↵shore at Ra's Al Hamra. Furthermore, the PGW presents various ranges of density varying over 1.5 kg.m 3 over di↵erent seasons, thus having PGW with di↵erent densities at the same locations.

"-What are the characteristics of the highly saline waters in the Sea of Oman and Arabian Sea, from a mesoscale and submesoscale perspective ?" The outflow appears fragmented, with small eddies, filaments and patches. In the Arabian Sea, the Persian Gulf Water is observed either as a coastal outflow in winter, ejected at Ra's Al Hadd in summer, or as a lens in spring. It can also be observed below the mesoscale eddies, being captured during their formations. This thesis, through the use of various dataset, in-situ measurements, satellite data, model reanalysis and primitive equations numerical simulations, studied the variability of the Arabian Sea and Sea of Oman over each season, from its surface circulation its impact on the water masses, through di↵erent scale, from the basin wide to the submesoscale. These circulations were also studied in term of their interannual and interseasonal variabilities, with both the observations and the model, with comparable results. Furthermore, process were studied with idealized models.

Perspectives

A further study beyond this thesis could be the creation of new climatologies using ARGO floats. During this study, the GDEM climatology was modified, using the anomaly between the measurements from the floats and the monthly temperature and salinity. In the region of interest it allowed us to have a better understanding of the mesoscale circulation. The GDEM climatology gave the best description of the northern Arabian Sea with numerous profiles extracted from MOODS each having the same weight, but not su cient to resolve the mesoscale. Adding the ARGO profiles to the climatology to study the mesoscale variability. The float and profile distribution was then su cient. This thesis could also lead to more process studies. Idealized models could allow us to study in particular the ejection of Rossby waves from the eastern coast, from their formation linked to a density front variation, to their interactions with the western eddies. As well idealized numerical simulations could study the process of deformation and advection of the Persian Gulf Water outflow.

Conclusion & Perspectives

An emphasis on the coastal dynamics, from the coast to 100 km o↵shore, with an impact of the topography would be of interest. The Phys-Indien 2014 experiment took place in March 2014 and recorded the thermohaline and velocity profiles above the shallow plateau o↵ Ra's Madrakah, along the southeastern coast of Oman. The mechanisms seem di↵erent from the o↵shore circulation, through stronger currents. The results of this experiment could also lead to the observation of submesoscale structures. Also very high resolution simulation (about 1 km horizontally) or intense observations over the Sea of Oman would allow to have a better understanding of the finer structure of the Persian Gulf outflow. These observations could be lead using SeaSoar, gliders and close-range XCTD to measure the thermohaline characteristics, VM-ADCP and seismic reflection of the upper water column would also lead to very high resolution of the velocity and density fields. This could be used to study the transport of Persian Gulf Water through the submesoscale structures, and their seasonality. As well, the mechanisms inducing the lateral injection of oceanic waters in the eddies, or above and below the the smaller structures could be of interest. For the surface fields, high resolution altimetry through SWOT, satellite surface temperature and surface winds via Synthetic Aperture Radar (SAR) would allow a remote sensing study of the mesoscale structure and the potential impact of the submesoscale on the surface.

Historically, the northwestern Indian Ocean has been of high interest through several experiments focusing on specific region or periods. Nowadays, an e↵ort is still put from the neighbouring countries, Saudi Arabia, Oman, Iran or India, but also from the United States and France (SHOM) to study this region. Numerical simulations are concentrating on the evaporative basins, the Red Sea and the Persian Gulf, but a few also focus on the Arabian Sea. As well, in-situ experiments, such as the French Phys-Indien program, show the interests of the regional dynamics and its impact on the biogeochemistry. Blooms of red plankton has been known to prevent the proper functioning of desalination plants in the region. This also supports e↵orts towards a better knowledge of the regional ocean dynamics, biogeochemistry and biology.

Introduction

In the Straits of Gibraltar, fresher north Atlantic water enters the Mediterranean Sea near the surface, and saltier Mediterranean Water flows westward near the bottom. This outflow of Mediterranean Water cascades down on the continental slope of the Gulf of Cadiz, mixes with surrounding north Atlantic Central Water, and veers northward under the influence of the Coriolis force. This Mediterranean Water (MW) forms several cores of mixed water, which stabilize at depths of 400, 800, and 1200 m (called, respectively, the coastal or shallow, upper, and lower cores of MW) [START_REF] Madelain | Influence de la topographie de fond sur l' ecoulement m editerran een entre le d etroit de Gibraltar et le cap Saint Vincent[END_REF][START_REF] Zenk | On the Mediterranean outflow west of Gibraltar, Meteor Forschungsergeb[END_REF][START_REF] Ambar | A shallow core of Mediterranean water off western Portugal[END_REF]. These cores follow the Iberian continental slope northwestward. Crossing the Portimão Canyon (near 8 30 0 W, 36 30 0 N) and then the Cape Saint Vincent canyon (near 9 30 0 W), the upper and lower cores of MW are destabilized and form mesoscale lenses, called MW cyclones and meddies [MW anticyclones; see in particular [START_REF] Armi | Large lenses of highly saline Mediterranean Water[END_REF][START_REF] Armi | Two years in the life of a Mediterranean salt lens[END_REF][START_REF] Prater | A meddy of Cape St Vincent. Part I: Description[END_REF][START_REF] Bower | Direct evidence of meddy formation off the southwestern coast of Portugal[END_REF][START_REF] Bower | Lagrangian observations of meddy formation during a Mediterranean undercurrent seeding experiment[END_REF][START_REF] Sadoux | Experimental study of the stability of an intermediate current and its interaction with a cape[END_REF][START_REF] Ch Erubin | Instability of the Mediterranean Water undercurrents southwest of Portugal: Effects of baroclinicity and of topography[END_REF][START_REF] Serra | Eddy generation in the Mediterranean undercurrent[END_REF][START_REF] Serra | Observations and numerical modeling of the Mediterranean Outflow splitting and eddy generation[END_REF][START_REF] Ambar | Observations of the Mediterranean Undercurrent and eddies in the Gulf of Cadiz during 2001[END_REF]. Several sites of meddy formation have been identified along the west coast of Portugal and Spain notably: the Estremadura Promontory [START_REF] K€ Ase | Observational evidence of salt lens formation in the Iberian basin[END_REF][START_REF] Zenk | New observations of meddy movement south of the Tejo plateau[END_REF] and Cape Ortegal [START_REF] Paillet | Real-time tracking of a northern meddy in 1997-98[END_REF][START_REF] Paillet | Dynamics and evolution of a northern meddy[END_REF].

Meddies are biconvex, anticyclonic lenses of MW with radii of 20-75 km, thicknesses of 500-1000 m, centered at 1000 m depth, and azimuthal velocities of 0.2-0.5 m s 21 [START_REF] Bower | Meddies, eddies, floats and boats: ''How do Mediterranean and Atlantic waters mix?[END_REF][START_REF] Richardson | A census of meddies tracked by floats[END_REF]. They contain warm and salty waters usually with a temperature maximum near 800 m depth and a salinity maximum near 1200 m depth. The average direction of meddy propagation, under the influence of beta effect, currents and topography, is southwestward [Colin de [START_REF] De Verdie `re | On the southward motion of Mediterranean salt lenses[END_REF][START_REF] Daniault | Circulation and mixing of the Mediterranean water west of the Iberian Peninsula[END_REF], but many deviations from this direction have been observed [K€ ase [START_REF] K€ Ase | Reconstructed Mediterranean salt lens trajectories[END_REF][START_REF] Shapiro | Spreading pattern and mesoscale structure of Mediterranean outflow in the Iberian Basin estimated from historical data[END_REF][START_REF] Richardson | Tracking three meddies with SOFAR floats[END_REF][START_REF] Richardson | A search for Meddies in historical data[END_REF][START_REF] Richardson | A census of meddies tracked by floats[END_REF][START_REF] Richardson | Meddy trajectories in the Canary basin measured during the SEMAPHORE experiment, 1993-1995[END_REF][START_REF] Tychensky | Hydrological and dynamical characterization of meddies in the Azores region: A paradigm for baroclinic vortex dynamics[END_REF][START_REF] Paillet | Real-time tracking of a northern meddy in 1997-98[END_REF][START_REF] Paillet | Dynamics and evolution of a northern meddy[END_REF][START_REF] Carton | Meddy dynamics and interaction with neighboring eddies southwest of Portugal: Observations and modeling[END_REF][START_REF] Barbosa Aguiar | A census of Meddies in a long-term high-resolution simulation[END_REF].

In [START_REF] Carton | Meddy dynamics and interaction with neighboring eddies southwest of Portugal: Observations and modeling[END_REF], the evolution of two meddies, detected and sampled southwest of Cape Saint Vincent in November 2000, was reported. These meddies interacted briefly and then followed different trajectories. The easternmost meddy coupled with a cyclone located on its southern side and drifted westward (as revealed by the analysis of altimetric maps). The westernmost meddy followed an unusual Y-shaped trajectory revealed by the tracking of Rafos floats seeded and trapped in this eddy. The first part of this trajectory resulted a priori from the coupling of the westernmost meddy with a cyclone initially located on its northern flank. The second part of this trajectory was interpreted as the coupling of this meddy with another cyclone coming from the northeastern Gulf of Cadiz.

In this paper, we analyze these eddies in detail. By joining two observational databases and satellite data, we have been able to identify the origin and structure of the latter cyclone: originally it was part of the dipole first located near Portimão Canyon. We describe the initial state of these dipoles, positions and hydrological characteristics, in November 2000 for the westernmost dipole, in February 2001 for the easternmost, Portimão, dipole, and then in June 2001 for the final dipole. These latter hydrological sections shed light on the origin of the water masses that this final dipole contained. We then compare the trajectories of the Rafos floats initially seeded in the two original meddies, providing evidence of a strong interaction between these two meddies. We also compare Sea Level Anomalies (SLA) and Sea Surface Temperature (SST) evolution with the trajectories of the floats. Finally, other Rafos float trajectories in the vicinity of these dipoles indicate the presence and evolution of other eddies in the Gulf of Cadiz at the same period. To confirm our hypothesis, we ran idealized simulations of meddy or deep eddy collisions. We compared their results with those of experimental studies of collisions of two dipoles [START_REF] Van Heijst | Dipole formation and collisions in a stratified fluid[END_REF]].

Data and Methods

Data Collection

The MedTop program, directed by the Centro de Oceanografia of the University of Lisbon, was designed to understand the role of bottom topography, and in particular of the Portimão Canyon, in the generation and further evolution of meddies and MW cyclones. This program carried out three experiments at sea in February, May, and September 2001. Five hydrological sections, with XBT casts and CTD stations closely spaced (five nautical miles, NM) on each section, were performed during each cruise (see Figure 1). Whenever such eddies were detected, Rafos floats were seeded in them. Results from the Meddy program can be found in [START_REF] Serra | Observations and numerical modeling of the Mediterranean Outflow splitting and eddy generation[END_REF], [START_REF] Ambar | Observations of the Mediterranean Undercurrent and eddies in the Gulf of Cadiz during 2001[END_REF], and [START_REF] Serra | Surface expression of Mediterranean Water dipoles and their contribution to the shelf/slope-Open ocean exchange[END_REF].

The Semane program, directed by the French Navy Oceanographic Service of Brest, was designed to understand the structure and variability of the MW cores and of detached eddies in the Gulf of Cadiz. This program carried out seven experiments at sea in November 1995, July 1997, July 1999, July 2000, November 2000, July 2001, and July 2002 (the latter experiment was carried out by IFREMER). Long CTD-LADCP sections across the Gulf were performed during these cruises and the results are described in other papers [START_REF] Quentel | Structure and temporal variability of Mediterranean Water from hydrological and marine seismic data south of Portimao Canyon (Gulf of Cadiz) from 1999 to 2002[END_REF][START_REF] Alves | Hydrological structure, circulation and water mass transport in the Gulf of Cadiz[END_REF]. These experiments also included XBT-XCTD cross sections of eddies, previously detected by altimetry or via other hydrological sections. Rafos floats, or deep-drogued Surdrift buoys, were also released in such eddies [see, for instance, Carton et al., 2002]. The XBT-XCTD casts of interest from Semane in November 2000 and June 2001 are shown on map (Figure 1).

For comparison with hydrological data, we used monthly mean fields of SLA obtained from AVISO with a 1/ 4 3 1/4 resolution on a Mercator grid and a daily interpolation. Informations and data sets can be retrieved from: http://www.aviso.altimetry.fr/en/data/products.html. Level 4, monthly averaged, SST maps with a 1/4 3 1/4 resolution were obtained from the Met Office Ostia system. Data can be retrieved from: http://www.ncof. co.uk/. SST anomalies were calculated from the 23 year climatology. Finally, topographic maps were obtained from the NGDC data center ETOPO1.

Data Processing

Hydrological profiles performed during these cruises measured temperature and salinity from surface to bottom (salinity was calibrated via chemical analysis of water samples conductivity and temperature). Visual analysis of each profile, and a median filter eliminated spurious points on the profiles. Then a binomial Newtonian smoothing was applied to reduce the number of points to 10 or 20% of the initial number, under the condition that the initial profile could be reconstructed from the final one with an error at most equal to the probe accuracy. With XBTs and XCTDs, the accuracy on temperature measurements is 0.018 C and that on salinity is 0.013 (the same accuracy is obtained for reconstructed salinity profiles). From a water mass diagram, salinity was reconstructed for XBTs using the method developed by K€ ase et al. [1996]. In Figure 1, we note that the XBT and XCTD casts were performed very closely. In order to calculate a regression of salinity on the locations of the XBT probes, we used the XTCD data. We plotted T-S diagram and inferred the properties of the MW and the surrounding waters. Those hydrological properties varying smoothly from one water mass to another, we could obtain salinity profiles associated with the temperature measured via XBT.

Many isobaric Rafos floats were released during the Semane and MedTop experiments, but we will only present here the trajectories of three floats. Float SV106 (renamed RF1 hereafter) was released in the northwestern meddy identified and sampled on 28 November 2000, at 35 36.40 0 N-10 19.19 0 W; this float was ballasted at 1000 dbar. Float me203 (RF2 hereafter) was released in a meddy at 36 08.46 0 N-8 55.45 0 W on 28 February 2001; it was ballasted at 1300 dbar. Float me206 (hereafter RF3) was released at 36 11.24 0 N-8 02.28 0 W on 23 February 2001; it was also ballasted at 1300 dbar. These floats were located by sound sources, a Portuguese source and two German sources (a distant source in the north Atlantic was used in case of ambiguities on the trajectories). Absolute precision of the positioning is 5 km, obtained as the distance between the last acoustic triangulation and the first surface location given by ARGOS, but relative errors within a trajectory are likely to be smaller.

Deep-drogued Surdrift buoys were released during Semane 2001. A 10 m long holeysock drogue was tethered to the surface buoy by a 2 mm thick Kevlar cable. These cables were 900-1200 m long. All Surdrift buoys were programmed for a 180 day mission and positioned via ARGOS. 1). Their thermohaline and dynamical characteristics, as well as their surface signatures will be described chronologically in the following sections. During the Semane 2000 experiment which took place in November 2000, two dipoles were identified by [START_REF] Carton | Meddy dynamics and interaction with neighboring eddies southwest of Portugal: Observations and modeling[END_REF] southwest of Cape Saint Vincent. The meddies were, respectively, located near 35 45 0 N, 10 75 0 W and 35 30 0 N, 10 15 0 W (see Figure 2a, for positions). They were accompanied by two cyclones detected on altimetry by pronounced depressions north and south of the meddies, respectively. For this study, we will focus on the northern dipole and refer to the meddy and the cyclone as M1 and C1. Furthermore, a Rafos float, RF1, was trapped in M1 for 9 months, 2b). We will call them M2 and C2, respectively. Two Rafos floats, RF2 and RF3 followed M2 from February to September and April to September, respectively.

Analysis of the Experiments Data Showing the Eddy Interaction

Finally, in June 2001, a third dipole was detected from Semane 2001 hydrological sections, with a meddy near 34 15 0 N, 9 30 0 W and a cyclone north of it at 35 N, 9 W (see Figure 2c). The three Rafos floats described earlier were found around this meddy and a Surdrift buoy (SF) was released inside it. Our assumption is that the first two dipoles (M1-C1 and M2-C2) collided and that the meddies merged to form M3. [from Carton et al., 2010]. The distance is measured from the beginning of the section at its southwesternmost point. This anticyclone has two cores, one with a temperature maximum 13.4 C near 700 m depth, and one with a salinity maximum of 36.8 near 1300 m. 3). This isocontour is chosen with respect to the maximum salinity and temperature gradient for this deeper and larger core of the meddy.

In [START_REF] Carton | Meddy dynamics and interaction with neighboring eddies southwest of Portugal: Observations and modeling[END_REF], cyclone C1 was found northeast of M1.

It was also shown that the interaction of M1 and C1 with another meddy and cyclone induced the coupling and eastward drift of M1-C1, as seen on RF1 trajectory (see Figure 4). After its eastward drift, float RF1 originally trapped in meddy M1 moved northward, then southward, and finally southeastward. Due to lack of in situ data at the time, [START_REF] Carton | Meddy dynamics and interaction with neighboring eddies southwest of Portugal: Observations and modeling[END_REF] used altimetric maps and a simple quasi-geostrophic model to show that this change of direction resulted a priori from the coupling of M1 with a cyclone other than C1.

Dipole Formed on Portimão Canyon in February 2001, M2-C2

MedTop hydrological sections, in February 2001, revealed the formation of a meddy coupled with a cyclone near Portimão Canyon, M2 and C2. Figure 5a shows the temperature and salinity distributions on the section from (36 14 0 N; 9 12 0 W) to (35 56 0 N; 8 14 0 W). In the northern part of this section lay meddy M2; this meddy occupied the 600-1500 m depth range. Its temperature maximum was 13.4 C near 900 m depth and its salinity reached 36.8 in a deep core between 1000 and 1400 m depths. Its radius, based on the 11.5 C isotherm, or on isohaline 36.2 psu, at 1200 m depth was about 25 km. The horizontal gradients of temperature and salinity at the periphery of meddy M2 were weaker toward northwest than southeast, where cyclone C2 lay.

Cyclone C2 had a temperature maximum equal to 11.4 C between 1100 and 1200 m depths; its salinity maximum, 36.4, was centered near 1200 m (see Figure 5a; Figure 5b provides another cross section of this cyclone). Furthermore, the spacing of isopynals (see Figure 6) became smaller between 1000 and 1400 m depth, characteristic of a cyclone; this was also confirmed by the calculation of the geostrophic velocity with a level of reference of no motion at 2000 m. Maximal velocities were on the order of 20.3 m s 21 for M2 and of 0.25 m s 21 for C2 and were found near 1200 m depth (see Figure 6). This figure also shows that:

1. The zero velocity line was tilted with respect to a vertical axis for both eddies; such a tilting is maybe achieved by the coupling between the two eddies in a stratified flow [START_REF] Dritschel | Vortex merger in rotating stratified flows[END_REF].

2. Both eddies had a surface signature in velocity, which was stronger above C2 than above M2; this intensification may be related to the existence of a separate structure above C2. The two velocity signals appear merged at 1200 m depth, but were separated by a third signal at the surface with an anticyclonic motion. This indicates that more dynamical structures (e.g., filaments) existed near the surface than at depth. Unfortunately, the map of Sea Level Anomaly for mid February 2001 (not shown) cannot evidence smallscale features at this location (due to the resolution of interpolated altimetric maps).

Considering the relative positions of M2 and C2, this dipole was moving southwestward at the time of its survey, in February 2001 (see Figure 7). After this strong interaction, the trajectory of RF2 veered southward. The second part of this trajectorybetween 35 30 0 N, 10 30 0 W and 34 N, 9 W seemed similar to that of RF1. To confirm this similarity, we separated the trajectories of floats RF1 and RF2 into mean drift and loops. The mean drift is represented in Figure 2c and was defined as the motion of the instantaneous center of rotation for each float. These were strikingly similar for both floats in this region. The center of rotation of RF2 was always less than 10 km away from that of RF1. Thus, the two floats must have drifted in the same meddy at that time. This meddy, M3, was the result of the interaction of M1 and M2.

We also show the trajectory of Rafos float RF3 (see Figure 7b). After its release east of Portimão Canyon, it followed the lower undercurrent of MW and drifted offshore along this canyon. It became trapped in an anticyclone (a meddy, near 36 15 0 N, 8 45 0 W) and was later ejected from this meddy to follow the continental slope. Again it veered offshore, from Cape Saint Vincent, and after two large anticyclonic loops, it reached the meddy containing RF1 and RF2, near 35 N, 10 W, in April 2001. This position also corresponded to the maximal temperature recorded by the float. From there on, its trajectory became comparable with that of RF2. The gradual decrease in temperature measured by the floats can be related to their progressive ejection from the meddy.

The Resulting Dipole in the Southern Gulf of Cadiz in July 2001, M3-C3

Figure 8 presents the zonal and meridional distributions of temperature and salinity of three selected XBT and XCTD sections, performed between 3 and 5 July, during Semane 2001 experiment. These sections were located between 34 N and 35 N meridionally, and 8 W and 10 W zonally.

On the quasi-meridional section (Figure 8a), one can identify a cyclone on the northern side and a meddy. The meddy position corresponds to the Rafos floats loops during this period. The cyclone is called C3 and the meddy M3 (labeled M6 in Figure 6 of [START_REF] Serra | Surface expression of Mediterranean Water dipoles and their contribution to the shelf/slope-Open ocean exchange[END_REF]). Now, we compare the hydrological properties of M3 with those of M1 and M2, and the characteristics of C3 with those of C2.

We recall that M1 and M2 had a 22 and 25 km radius, both based on isotherm 11.5 C, or isohaline 36.2 psu. Figure 8b indicates a radius of about 40 km for meddy M3, with the same isoline. Therefore, the salt and heat budgets advected in M1 and M2 cores, from isohaline 36.3 psu to focus on the cores of the meddies and to avoid diffusion, were, respectively, 3.4 3 10 13 and 3.7 3 10 13 kg of salt, and 4.5 3 10 19 and 4.8 3 10 19 J. The salt quantity calculated for M3 was 7.7 3 10 13 kg and the heat was above 9 3 10 19 J. The salt and heat from the inner cores of M1 and M2 almost equal those of M3, which indicates that an important mass transfer must have occurred from M2 to M1 to form M3. In terms of volume, M1 and M2 had vertical extensions of 800-900 and 650-700 m, respectively, and M3 about 950-1000 m. With respect to their radii, M3 had a more important volume (about 5 3 10 12 m 3 ) than both M1 and M2 together (about 1.2 3 10 12 and 1.4 3 10 12 m 3 ), this could be due to entrainment of surrounding waters.

Figure 8b shows the salinity anomaly of M3 located between 600 and 1500 m depths, with a double core, similar to that of M1, and slightly more surface intensified than that of M2. On the Brunt-V€ ais€ al€ a frequency (not shown) two local minima were found between 500-800 and 1000-1500 m showing stability in those cores. This tendency was also found on M1 but M2 had only a local minimum between 1000 and 1500 m. The salinity peak of M3 (36.6 psu) was slightly below those of M1 and M2 (36.8 psu). The temperature of M3 reached 13.5 C and its anomaly was marked between 600 and 800 m depths, being closer to that of M1 than to that of M2.

Our interpretation is that during the merging process, meddy M2 wrapped around meddy M1, before their waters became mixed, transferring heat and salt to M1. The observed turbulent motions are likely to be related also to the kinetic energy and angular momentum redistributions caused by the meddy merger. During this merging, entrainment of surrounding North Atlantic Central Water (NACW) and eventually traces of North Atlantic Deep Water (NADW) led to a slight decrease in salinity.

Figure 9 shows the horizontal maps of temperature and salinity on three different isopycnals, obtained by optimal interpolation of the data over the 3 days of measurements, and using a 15 km radius of correlation; this radius corresponds to the average distance between two measurement points and the desired resolution but is smaller than that commonly used in the open ocean [START_REF] Arhan | Dynamics of eddy motions in the eastern north Atlantic[END_REF]. This is L'H EGARET ET AL.

V C 2014. American Geophysical Union. All Rights Reserved. 10 7.5. Analysis of the experiments data showing the eddy interaction also explained by the fact that we study mesoscale motions in a region with a moderately small deformation radius. At r 1 5 31.9 kg m 23 , isopycnal with a mean depth of 800 m, the warm waters of the cyclone (north) and meddy (south), with temperatures on the order of 12-12.5 C were separated by a strip of colder water (on the order of 10 C); this strip was about 30 km wide meridionally. The salinity distribution displayed a similar pattern, with two peaks at 36.2 and 36.4 psu (for the meddy and cyclone) separated by a strip of freshwater (with salinity 35.6-35.7 psu). This water is NACW. Supposedly, this strip has been advected eastward by the joint rotation of the meddy and cyclone; it has been elongated by the strain field created by the two eddies. This situation is reminiscent of that observed during Semane 99 [Carton et al., 2002] but here without an external source of Mediterranean Water.

The isopycnal r 1 5 32 kg m 23 , having a mean depth of 1000 m, is situated between the two cores of the eddies, resulting in lower values of temperature and salinity for the meddy and the cyclone (12.5 C and 36.5 psu for the meddy versus 11.5 C and 36.2 psu for the cyclone). At r 1 5 32.2 kg m 23 , isopycnal with a mean depth of 1200 m, the temperature of the meddy was much higher than that of the cyclone (12 C versus 10. 5-11 C). This could be due to the entrainment of NADW below the cyclone core during its formation, and to its mixing with Mediterranean Water. Finally, for the same isopycnal, the water tongue between the two eddies may have resulted from a turbulent horizontal diffusion of salt due to the strong interaction between these eddies.

The geostrophic velocity maximum calculated in M3 (not shown), with a level of no motion at 2000 m, was about 0.25 m s 21 at 750 m depth and at a radius of about 10 km from the meddy center. The velocity was still 0.15 m s 21 at 1000 m depth. These values were slightly below the velocities recorded by the float (about 0.2 m s 21 ). The M1 and M2 velocities were, respectively, 0.3 and 0.5 m s 21 . Therefore, both the interaction and propagation off the eddies have affected their kinetic energies. With respect to their volume, the sum of M1 and M2 kinetic energies (about 6 3 10 10 J) was more important than the M3 kinetic energy (about 3.8 3 10 10 J), here again it could be due to entrainment of surrounding waters around M3, and/or to the loss of M1 or M2 fluid to the environment. Kinetic energy variation after eddy merger has been discussed in theoretical studies (see below, the modeling part). L'H EGARET ET AL.

V C 2014. American Geophysical Union. All Rights Reserved.

12 Journal of Geophysical Research: Oceans To understand more precisely the interaction between M1 and M2 (with the influence of C1 and C2), we segmented in Figure 10 the trajectories of floats RF1, RF2, and RF3, in monthly periods from March to August. In July and August, float RF1 is substituted by float SF released in M3. We recall that SF is a Surdrift buoy launched in M3 in June 2001. We superimposed these segments on the SLA map corresponding to the middle of each period. We named the main features SA and SC for surface anticyclone and surface cyclone. We also compared the trajectories of the floats with SST anomaly that is the difference between a monthly averaged map of SST and 23 year climatological months. We repeated the positions and names of the main features detected on altimetry on these SST anomaly maps. The joint use of SLA and SST anomalies gave us an indication of mid depth structures [START_REF] Bashmachnikov | Manifestation of two meddies in altimetry and sea-surface temperature[END_REF]. Positive SLA and SST anomalies show that an anticyclone is located near the surface. Negative anomalies for both altimetry and SST correspond to a cyclone. Positive SLA anomaly with a negative SST anomaly corresponds to an eddy that lies at depth, e.g., a meddy.

On the first map (March 2001, Figure 10a), all floats were clearly separated. RF1, released in M1 was looping anticyclonically under a patch of positive SLA (SA2 on the maps) near 36 N, 10 W; RF2, released in M2 was rotating anticyclonically with small loops (20 km in diameter) while following a clockwise trajectory around the aforementioned positive SLA. RF3 was moving almost zonally on the southern side of the Iberian peninsula coast but it headed southward along 8 45 0 W and started looping anticyclonically (with small loops, of 15 km diameter) under a negative patch of SLA, which could be continuously tracked until the Semane 2001 survey in July. We call SC2 the cyclone corresponding to this negative SLA. Furthermore, at that time, RF3 was quite far (by about 80 km) from the two other floats. This may indicate that it was trapped in a small anticyclonic eddy of MW formed after M2 or as a spin-off of M2 formation. In Figure 11a, the corresponding SST anomaly field showed a negative anomaly advected southward, at the position of SC2 (36 N and 9 30 0 W). This can be the signature of a subsurface cyclone, C3.

In April, the correlation between trajectories and altimetry becomes clearer (see Figure 10b), we observe here the merger. RF1 and RF2 loops became much larger (80 km in diameter, suggesting a merger of M1 and M2). Both floats rotated anticyclonically under a positive SLA patch (SA2) extending from 35 N to 37 N along 10 30 0 W. This positive SLA matched in position to a strongly SST negative anomaly corresponding to a meddy (Figure 11b). A substantial amount of fluid must have been transferred from M2 onto M1 at that time explaining the exchange of float RF2 between the two meddies and leading to the final meddy M3. This negative SST anomaly patch also extends toward the altimetric signature of C3, suggesting that M3 was elongated due to shear created by the neighboring cyclone (C3). At the same time, float RF3 circled cyclonically around SC2 near 36 30 0 N, 9 30 0 W. The trajectory of this float also displayed two anticyclonic meanders a priori influenced by meddy M3.

In May, all three floats started drifting in M3 whose altimetric signature gradually decreased in amplitude, while the altimetric signature of the accompanying cyclone (C3) remained strong. The SST negative anomaly from both SC2 and SA2 intensified. The diameter of the RF2 loops decreased from April to May, showing an adjustment of M3 after the merger of M1 and M2. The temperature measured by RF2 also dropped by 1 C at that time (see Figure 7a). This can be explained as follows: the M1 and M2 interaction increased the thermohaline gradients of M3; thus, any radial shift of float trajectory in the meddy may have led to a noticeable increase or decrease in the measured temperature.

In June 2001, the loops in the trajectories of the three floats became elongated; as the meddy M3 became deformed elliptically. This deformation is coherent with an anticlockwise acceleration of M3 around C3 at that time. Float RF2 recorded a lower temperature as it moved outward in the meddy, starting in early July 2001.

These float trajectories show that M3 was following a cyclonic trajectory in the Gulf of Cadiz, a priori under the influence of the neighboring vortices (and in particular of C3), but this trajectory was also in agreement with the regional circulation. The SST anomalies map of June (Figure 11d) also showed a cyclonic motion of the warmer water near M3 and around C3.

During Semane 2001, several deep-drogued Surdrift buoys were used. Here, we study in particular SF which was released on 5 July 2001 at 34 12 0 N, 9 05 0 W, and had a drogue at 1050 m depth. This buoy performed several loops in M3 and drifted in a northeastward direction before turning back southwestward along the Moroccan coast.

L'H EGARET ET AL.

V C 2014. American Geophysical Union. All Rights Reserved.

13

Journal of Geophysical Research: Oceans L'H EGARET ET AL.

V C 2014. American Geophysical Union. All Rights Reserved.

In July, all floats and buoys were looping in M3, which was heading northeastward (see Figure 10e). In August, the surface signature of M3 intensified while this meddy was still rotating around C3. This signature became also strongly elongated at that time, as if M3 was joining another eddy; the SST anomalies map (Figure 11f) showed a propagation of warm water near 36 N-8 30 0 W). However, the trajectories of Rafos floats RF2 and RF3 do not show any sign of elongation [START_REF] Lamb | Hydrodynamics[END_REF]. Note that the drogue of buoy SF broke around that time, and that the buoy then performed small inertial loops while drifting southwestward. A multicycle acoustically tracked float from the ARCANE experiment, drifting at 450 dbar, performed anticyclonic loops and drifted anticlockwise in the Gulf of Cadiz during July, August, and September 2001 around C3. Its trajectory is shown in Figure 12. We are primarily interested in the segment of trajectory in July 2001. There is a clear correspondence between the float position on July and the SLA map with a positive patch near 36 N, 8 W. Thus, there was another anticyclone (SA4, not shown on the map), perhaps a meddy, in the eastern Gulf of Cadiz during the same lapse of time. Finally, at the end of August, buoy SF was ejected from M3 while RF2 and RF3 were still regularly rotating in this eddy, without any influence of a meddy to its north. During July and August, the surface signature of C3 barely evolved.

Numerical Modeling of the MW Dipole Collision

The Model

We used a three-layer quasi-geostrophic model to investigate the mechanism of interaction between dipoles M1-C1 and M2-C2. Since this interaction took place far from the continents and high above the bottom, we used a flat-bottom, open-ocean model. More specifically, the model domain is square and horizontally biperiodic, but the eddy interaction takes place near the center of the domain so that periodicity influences are negligible. A beta-plane configuration was implemented.

We chose a length scale L 5 40 km and a velocity scale V 0 5 1 m s 21 . Our dimensionless domain length was 4p or in dimensional terms 502 km. This is about 4.5 of latitude and 5.5 of longitude. The resolution was 256 3 256 points (one point every 2 km).

The quasi-geostrophic equations for potential vorticity are in layerwise form: dq j dt 5 m 2 r 4 w j 1m 4 r 6 w j where w j ðx; y; tÞ is the stream function and q j is the potential vorticity in layer j q j 5r 2 w j 1F j;j21 ðw j21 2w j Þ1F j;j11 ðw j11 2w j Þ1f 0 1by L'H EGARET ET AL.

V C 2014. American Geophysical Union. All Rights Reserved.

18 7.6. Numerical modeling of the MW dipole collision is frequent with asymmetric merger, a fragment of M2 is shed during the process (in the observations, it seems that all of M2 merged with M1). After 5 weeks, the two meddies have separated and about 30% of M2 has remained coupled with C1. This asymmetric dipole will rotate cyclonically over a short distance in the northern part of the domain as can be seen on potential vorticity plots of weeks 7 and 20. Thus, there has been an exchange of cyclones between the two meddies. On the same plots, it can be seen that C2 and what is now called M3, drift first southwestward and then southeastward, also cyclonically, as in the observations. This suggests that C3 is C2.

The efficiency of this merger process between M1 and M2 is now quantified with two diagnostics.

First, we seeded artificial floats in each meddy, at the initial stage of the simulation. These floats were seeded by groups of 10, regularly spaced azimuthally, and at different distances D f from each meddy center. We let D f range from 0.1 to 0.7 meddy radii. In Figure 14, we show how many floats initially in M2 were finally found in M1 (''M1gainedM2'') and how many floats initially in M1 were lost to the surrounding flow (''M1lost''). This figure shows that floats were torn from deep in the M2 core (10% of the flows at D f 50:25R were finally found in M1). This figure increases to 80% for D f 50:4R and it decreases to 50% for D f 50:67R. This is simply explained by the fact that peripheral floats can be expelled from M2 into the surrounding ocean (e.g., in filaments) and not be absorbed by M1. The second curve shows that M1 itself lost its peripheral floats during the merging process.

In short, M2 was severely disrupted by the merging process (as far inward as a quarter of its radius) and lost much of its mass essentially to M1 but also to the surrounding ocean. This is compatible with the fact that only a minor fraction of M2 finally drifted with C1. During merger, M1 also lost a part of its external envelope.

The second diagnostics that we performed was the computation of the area integral of potential vorticity for M1 in its main (intermediate) layer. We found that M1 gained 50% of surface integrated potential vorticity. Again, such a value is compatible with the final remainder of M2. This value can be also used as an estimate for the change in tracer contents (e.g., salt contents). Theoretical studies [e.g., Cushman Roisin, 1989] have detailed the conditions under which the conservation of local or integral invariants is possible during eddy merger. In particular, the conservation of both potential vorticity, angular momentum, mass and kinetic energy, was shown possible only if filaments were formed at the end of the merging process. These filaments (modeled as a vorticity ring in the theory) would carry only a minor fraction of mass or of kinetic energy but the vast majority of angular momentum. This has two implications: first, these filaments are usually thin and are more sensitive to dissipation so that part of the initial kinetic energy can be lost after merger; second, as stated in Cushman Roisin [1989], the process of eddy growth and of filament formation L'H EGARET ET AL.

V C 2014. American Geophysical Union. All Rights Reserved.

19

Chapter7. Evidence of Mediterranean Water dipole collision in the Gulf of Cadiz is the physical manifestation of the inverse energy and direct enstrophy cascades, typical of quasi-2-D turbulence.

4.3. Sensitivity Tests and Constraints for the Realization of the Observed Phenomenon It must be noted that this evolution also occurred for other initial conditions. First, if the eddies had slightly different velocities or positions, partial merger of M1 and M2 was also observed. But the efficiency of the merger process was weaker in the other experiments we performed. Also, the final trajectory of the stronger dipole was more often southwestward.

Second, even if M1-C1 drifted eastward from an initial western position, merger of M1 with M2 could be obtained with a further cyclonic trajectory of the resulting dipole (southwestward, then southeastward; plots not shown here). Thus, this evolution is not unique. Therefore, we search the conditions under which this merger and/or the exchange of cyclonic companion can occur. To this aim, many numerical simulations were conducted.

As a preliminary note, we can recall that it is impossible to forecast exactly the evolution of the four vortex system due to the irreversible processes which occur during the interaction (vortex deformation, filamentation, merger, or splitting). This interaction is not comparable with that of four point vortices, except if they are widely separated (furthermore, the evolution of four point vortices is most often chaotic).

Collision of Collinear, Opposite, Symmetric Dipoles

To assess how the offset collision of dipoles can result in the merger of two of their eddies, we first analyzed the collision of two equal, symmetric dipoles on the beta plane. These dipoles were propagating in opposite directions (one eastward, the other one westward). We varied the offset between their (zonal) propagation axes (see Table 3). Hereafter, we call A the anticyclones and C the cyclones, which may not have the characteristics of the meddies nor of the MW cyclones.

Whatever the case (head-on or offset collision), no merger occurred between the anticyclones, even if they partially interacted for a short time. This temporary interaction and exchange of fluid is in agreement with the study by Velasco Fuentes and [START_REF] Velasco Fuentes | Collision of dipolar vortices on a b plane[END_REF]. Here, we observed that the cyclones simply exchanged partners and finally rotated to adopt an eastward propagation on the beta plane (the stationary configuration) [see Velasco [START_REF] Velasco Fuentes | Collision of dipolar vortices on a b plane[END_REF]. Changing the intensity or size of the eddies (to 0.5 or to 0.7) did not change this evolution qualitatively.

Collision of Opposite, Symmetric Dipoles With a Finite Angle Between Orientations

Starting from this observation, we studied the collision of the same symmetric dipoles, but now propagating southeastward and southwestward, respectively (Table 4). This initial orientation should favor the merger of the anticyclones, provided that they are located at the proper initial position.

For A1-C1, we chose an initial position symmetric of that of M2-C2 in the observations. For A2-C2, we allowed a variation in latitude to create an offset in the collision. These initial positions were calculated with the help of a point vortex model to obtain a first guess of the dipole trajectories.

Only when the two anticyclones came close to each other in the collision, did they merge (for an offset of about 0.6); under other circumstances, the shear and strain created by the neighboring cyclones was sufficient to pull the anticyclones apart. The most efficient merger occurred if A2 passed 5 km south of A1 (offset of 0.7). In that case, the merger efficiency was about 90%. This efficiency decreased to 50% if the anticyclones drifted 10 km apart (offset of 0.8).

Collision of Opposite, Asymmetric Dipoles

Finally, we studied the collision of two opposite, asymmetric dipoles. The ratio of stronger to weaker eddy velocity in each dipole was 0.6 (the dipoles were initially NW and NE of the domain; see Table 5). To find the 7.6. Numerical modeling of the MW dipole collision initial position of the dipoles, a four point vortex model was used, with equivalent circulations. Though this point vortex model was on the f-plane, it provided a fair estimate of the initial trajectory of the dipoles (before they interacted). This way, we could control the offset of the dipoles when they collided.

In this case, strong interaction between the anticyclones took place when they collided head-on (offset of about 0.6). But since C1 was stronger than C2, it paired with the strongest anticyclone resulting from the interaction. Only the weakest anticyclone drifted southwestward with C2. Again the most efficient merger occurred if A2 passed 5 km south of A1. In that case, the merger efficiency was about 40%.

Conclusions of the Modeling Part

This very simple modeling of the collision between two dipoles on the beta plane has shown that the observed situation could be reproduced with few parameters (eddy radii, intensities, positions) for the model and a simplified structure (Gaussian distribution of stream function) for the eddies. Setting the eddy parameters on the available observations, the observed meddy merger and further trajectory could be reproduced fairly well; in the model, meddy merger was not complete, contrary to observations, but the strength of all eddies involved was not exactly known at the time; this may account for this difference. Several diagnostics were implemented on this simulation, indicating that M2 was strongly eroded during the merging process. A substantial part of its mass was absorbed by M1 (as shown by the float count and by the integral budget of potential vorticity). Other experiments showed that meddy merger and southward drift also occurred when the eddy parameters were slightly changed or if M1-C1 came from the west instead of the north.

A series of idealized experiments, designed to test the conditions under which this evolution could take place, provided the following results:

1. the head-on or offset collision of collinear symmetric dipoles cannot lead to the same evolution because of the shear created by the cyclones on the anticyclones; this shear separates the anticyclones after they start merging; this is in agreement with a previous study by Velasco [START_REF] Velasco Fuentes | Collision of dipolar vortices on a b plane[END_REF];

2. the head-on or offset collision of asymmetric dipoles coming from the NW and from the NE can lead to anticyclone merger but the stronger resulting anticyclone then moves eastward with the stronger cyclone; 3. the lateral collision (at a finite angle) of such dipoles can lead to a merger of the anticyclones if these latter collide nearly head-on, but again this merger is not complete. Thus, the complete merger of M1 and M2 was certainly due to their position during the collision (only slightly shifted) and to the fact that the accompanying cyclones were not strong enough to tear apart the resulting meddy, but sufficiently strong (and comparable in strengths) to advect it anticlockwise after the merger.

Conclusion and Discussion

A unique joint data set has allowed the characterization of two dipoles of Mediterranean Water in the Gulf of Cadiz. The western dipole (M1 and C1) could not be completely described by lack of hydrological data, but its meddy was fully studied in [START_REF] Carton | Meddy dynamics and interaction with neighboring eddies southwest of Portugal: Observations and modeling[END_REF]. The Portimão dipole (M2 and C2), which had been described previously in [START_REF] Serra | Observations and numerical modeling of the Mediterranean Outflow splitting and eddy generation[END_REF], [START_REF] Ambar | Observations of the Mediterranean Undercurrent and eddies in the Gulf of Cadiz during 2001[END_REF][START_REF] Serra | Surface expression of Mediterranean Water dipoles and their contribution to the shelf/slope-Open ocean exchange[END_REF], has been identified as the structure colliding with the western dipole, in the present paper.

The western dipole drifted northeastward from its position southwest of Cape Saint Vincent, while the Portimão dipole drifted southwestward from the Canyon. They collided near 35 30 0 N, 10 W. In this process, the two meddies merged while the cyclones drifted apart. More precisely, M2 wrapped around M1 before merging with it and forming M3. M3 then drifted southeastward, coupled with a cyclone that we called C3. It should be noted that only a few complete observations of oceanic eddy merger are reported in the literature [START_REF] Creswell | The coalescence of two East-Australian current warm-core eddies[END_REF][START_REF] Schultz-Tokos | Merging and migration of two meddies[END_REF][START_REF] Richardson | A census of meddies tracked by floats[END_REF]. They showed mass

21

Chapter7. Evidence of Mediterranean Water dipole collision in the Gulf of Cadiz transfer between meddies. This collision is original because it presented the interaction of two dipoles, instead of meddies.

Eddies M3 and C3 were finally surveyed in early July 2001 in the southern Gulf of Cadiz during a Semane experiment. These hydrological sections confirmed that cyclone C3 was in fact C2 and that M3 was much larger than M2 or M1, thus supporting the hypothesis of a large mass transfer between M2 and M1. The similarity of the thermohaline characteristics also supported our description of the dipole interaction.

In fact, the hydrological and dynamical situation in the Gulf of Cadiz in April 2001 to June 2001 was more complex than the interaction of two dipoles of MW: surface maps and float trajectories point toward a very complex mesoscale field, a priori with many meddies and cyclones in the Gulf of Cadiz (as suggested by the altimetric maps of Figure 10). The general motion of these meddies and cyclones was anticlockwise east of 11 W.

More generally, this shows that meddy collisions/mergers are likely to occur close to sites of meddy (and MW cyclone) generation. The meddy-cyclone coupling, often observed recently, leads to a faster propagation velocity than that of the beta-drift of a meddy, and also to more energetic and more complex collisions (larger ''efficient section''). As such, these dipoles collisions, at the MW level, may be an efficient mechanism to generate the larger meddies that are encountered further outside of the Gulf of Cadiz [e.g., [START_REF] Armi | Large lenses of highly saline Mediterranean Water[END_REF]. This would make the Gulf of Cadiz a ''cauldron,'' somehow alike to the ''Cape Cauldron'' [START_REF] Boebel | The Cape Cauldron: A regime of turbulent inter-ocean exchange[END_REF]. These strong nonlinear interactions may also trigger spiral arms recently observed around meddies [START_REF] Song | Meddy, spiral arms, and mixing mechanisms viewed by seismic imaging in the Tagus Abyssal Plain (SW Iberia)[END_REF][START_REF] Hua | Arms winding around a Meddy seen in seismic reflection data south of the Azores front[END_REF].

Geophysical fluid dynamics states that vortex dipoles steadily propagate eastward on the beta plane while single eddies propagate westward. Thus, dipoles that are generated near Portimão Canyon rotate to adopt an eastward-propagating configuration. This may explain the anticlockwise general motion in the Gulf of Cadiz. Once the dipoles are blocked against the topographic step at 8 W, they tend to rotate again or to split [see also Carton et al., 2002].

Again, this raises the question of the origin of the cyclonic circulation in the Gulf of Cadiz: is it due to the influence of the water mass exchange at Gibraltar [START_REF] Peliz | Surface circulation in the Gulf of Cadiz. Part 2: Inflow/outflow coupling and the Gulf of Cadiz Slope Current[END_REF][START_REF] Lamas | Evidence of time-mean cyclonic cell southwest of Iberian Peninsula: The Mediterranean Outflow-driven beta-plume?[END_REF], to that of the general circulation, or to the eddy field [START_REF] Alves | Hydrological structure, circulation and water mass transport in the Gulf of Cadiz[END_REF] this study]? In fact, all these studies point toward a strong three-dimensional coupling of motions in the Gulf of Cadiz. Presently, this coupling cannot be easily quantified because existing data sets lack synopticity and duration, and because numerical models do not yet fully reproduce the turbulent field in the Gulf. More complete data sets, perhaps relying on a large number of profiling floats and gliders, and very high-resolution numerical models, should be able to quantify more precisely the role of each effect in this cyclonic circulation and in its variability.

  Ensuite je remercie mes amis pour m'avoir soutenu durant toutes ces années, et la liste est longue. Je remercie JK et Yann pour avoir toujours été à mes côtés. Merci à Nico, mon adversaire de toujours en géographie et compagnon de voyage. Merci à Loïc de me forcer à a ner mon sens critique et à Antoine d'être toujours à l'écoute. Merci à Thomas et Laurence pour leurs connaissances a↵utées et leurs volontés de protéger les animaux et la nature. Merci à Brendan (Tanguy Doudou) pour sa bonne humeur communicative et ses imitations de Arnaud M. Merci à Liliane pour avoir rempli mon estomac de délicieux mets sucrés. Merci à Gwen et Carole d'avoir bien élevé Loïc et JK et pour leurs bonnes humeurs. Merci à Dani d'avoir pris ma succession chez les Lyons. Merci à Julia pour ses conseils salvateurs avant la soutenance. Merci à Julien d'avoir corrigé mes goûts musicaux et culturels. Merci à Benjamin d'être mon ami depuis si longtemps. Merci à Suzanne et Yso de m'avoir confié Ilwë. Merci d'ailleurs à Boodhy qui m'a fait prendre souvent l'air durant les derniers mois de ma thèse. Merci au trio dynamique Maxime, Colin et Simon d'avoir repris le flabeau de l'AssESciB. Merci à Joss, Aurélia et Camila, mes bons

Figure 1 :

 1 Figure 1: Topographic map of the northern Indian Ocean.

  . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13 1.1.1 Topography . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13 1.1.2 Hydrological basins . . . . . . . . . . . . . . . . . . . . . . . . . . 16 1.2 Climate . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17

  1.2.1 Arabian Sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17 1.2.2 Gulf of Aden and Red Sea . . . . . . . . . . . . . . . . . . . . . . 18 1.2.3 Sea of Oman and Persian Gulf . . . . . . . . . . . . . . . . . . . 19 1.3 Oceanographic exploration, Historical discoveries and experiments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 1.4 Water mass formation and circulation . . . . . . . . . . . . . . . 20 1.4.1 Red Sea and Gulf of Aden . . . . . . . . . . . . . . . . . . . . . . 21 1.4.2 Persian Gulf and Sea of Oman . . . . . . . . . . . . . . . . . . . 23 1.4.3 Arabian Sea . . . . . . . . . . . . . . . . . . . . . . . . . . . .

Figure 1 . 1 :

 11 Figure 1.1: Map of the bathymetry and orography of the region of interest, clear plate frontiers are underlined

  Figure 1.3: ⇥/S diagram extracted from the GDEM climatology, averaged over a month for each monsoon, in July (blue) and January (red) for the Red Sea (top) and the Gulf of Aden (bottom). Abbreviations are : RSSW, Red Sea Surface Water; RSDW, Red Sea Deep Water; RSOW, Red Sea Overflow Water; GAIW, Gulf of Aden Intermediate Water and IOCW, Indian Ocean Central Water.

Figure 1 . 4 :

 14 Figure 1.4: ⇥/S diagram extracted from the GDEM climatology, averaged over a month for each monsoon, in July (blue) and January (red) for the Persian Gulf (top) and the Sea of Oman (bottom). Abbreviations are : ASHSW, Arabian Sea High Salinity Water; PGW, Persian Gul Water Water; IOCW, Indian Ocean Central Water, IODW, Indian Ocean Deep Water and AABW, Antarctic Bottom Water.

  Figure 1.5: ⇥/S diagram extracted from the GDEM climatology, averaged over a month for each monsoon, in July (blue) and January (red) for the Arabian Sea, West of 65 o E (top) and East of 65 o E (bottom). Abbreviations are : ASHSW, Arabian Sea High Salinity Water; PGW, Persian Gul Water Water; BBW, Bay of Bengal Water; IOCW, Indian Ocean Central Water, IODW, Indian Ocean Deep Water and AABW, Antarctic Bottom Water.

Figure 2 . 1 :

 21 Figure 2.1: Map with the measurement locations from Phys-Indien superimposed. The red line follows the VM-ADCP measurements and SeaSoar (with the exception of the Red Sea and Gulf of Aden), XBT-XCTD (blue crosses) and CTD (green circles). The black crosses represent the first positions of the float (Surdrift and PROVOR) launched during the experiment.

Figure 2 . 2 :

 22 Figure 2.2: Section of the SeaSoar without any correction (up) and with correction coecients applied (bottom).

Figure 2 . 4 :

 24 Figure 2.4: Section of VM-ADCP section along the coast of Oman without any filtering.

Figure 2 . 5 :

 25 Figure 2.5: Section of VM-ADCP section along the coast of Oman with the erronaous values filtered (up) and with a low-pass filter applied (bottom).

Figure 2 . 6 :

 26 Figure 2.6: Maps of the climatological salinity and temperature from GDEM at 200 m depth (up); and corrected via the ARGO floats averaged over one month, August (bottom).

Figure 2 . 7 :Figure 2 . 8 :

 2728 Figure 2.7: Map of the associated distribution of the ARGO floats and their representativity.

Figure 2 . 9 :

 29 Figure2.9: Maps of the MADT anomaly of the sea surface computed from the SSH signal (up) and the sea level anomaly from AVISO (right).

Figure 1 : 5 Fig. 1 .

 151 Figure 1: Bathymetry of the Arabian Sea (from ETOPO1 database)

9 Fig. 2 .

 92 Fig. 2. Trajectory of float 1900432 in the Gulf of Aden, superimposed on bathymetry; the green dot is the starting point of the trajectory; positions every 10 days are indicated by a red dot.

Figure 3 :Fig. 3 .

 33 Figure 3: Distribution of Lagrangian kinetic energy (in m 2 /s 2 ) with respect to horizontal displacement (top) and to the orientation of the displacement (bottom), for float 1900432 10 Fig. 3. Distribution of Lagrangian kinetic energy (in m 2 s 2 ) with respect to horizontal displacement (top) and to the orientation of the displacement (bottom), for float 1900432.

Figure 4 :

 4 Figure 4: Sea-level anomaly (SLA) map of May 16th and June 27th 2007 (days 40 and 80 of the mission -from top to bottom, previous page); SLA map of July 25th 2007 (day 108 of the mission -this page). The trajectory of float 1900432 is superimposed on each map. Red dots correspond to the position of this float every 4 days; the number close to the dots is the mission day shown only every 8 days for clarity.

Fig. 4 .

 4 Fig. 4. From top to bottom, sea-level anomaly (SLA) maps of 16 May, 27 June and 25 July 2007 (days 40, 80 and 108 of the mission). The trajectory of float 1900432 is superimposed on each map. Red dots correspond to the position of this float every 4 days; the number close to the dots is the mission day shown only every 8 days for clarity.

FloatFigure 5 :

 5 Figure 5: (Left) Vertical section of temperature (in o C) along the trajectory of float 1900432 in the Gulf of Aden; (right) vertical section of salinity

Fig. 5 .Figure 6 :Fig. 6 .

 566 Fig. 5. (Left) Vertical section of temperature (in C) along the trajectory of float 1900432 in the Gulf of Aden; (right) vertical section of salinity.

  Sci., 8, 227-248, 2012 www.ocean-sci.net/8/227/2012/ Chapter3. Mesoscale variability of water masses in the Arabian Sea as revealed by Argo floats

Figure 7 :Fig. 7 .

 77 Figure 7: Distribution of Lagrangian kinetic energy (in m 2 /s 2 ) with respect to horizontal displacement (top) and to the orientation of the flow (bottom), for float 2900392 17 Fig. 7. Distribution of Lagrangian kinetic energy (in m 2 s 2 ) with respect to horizontal displacement (top) and to the orientation of the flow (bottom), for float 2900392.

Figure 8 :Fig. 8 .Figure 8 :Figure 9 :

 8889 Figure 8: Sea-level anomaly (SLA) maps of January 25th and March 8th 2006 (days 115 and 160 of the mission -from top to bottom, previous page); SLA map of August 23th, 2006 (day 325 of the mission -this page). The trajectory of float 2900392 is superimposed. Dots indicate the position of the float every 5 days. Mission days are indicated next to the dots.

  19

Fig. 9 .

 9 Fig. 9. Vertical sections of temperature (in C, left) and of salinity (right) along float 2900392 trajectory for year 1.

Figure 10 :Fig. 10 .

 1010 Figure10: Trajectory of float 2900392 in the Gulf of Aden, superimposed on bathymetry, year 2; the green dot is the starting point of the trajectory; positions every 20 days are indicated by a red dot.

  released on 14 March 2006 at 54 02 0 E, 11 42 N. Its last data were collected 25 August 2009 at 60 45 0 E, 17 10 0 N. Here, we present the first two years of its evolution. Its cycle period is 10 days.

Figure 11 :

 11 Figure 11: Sea-level anomaly (SLA) maps of November 22nd, 2006 and January 17th, 2007 (days 420 and 475 of the mission -from top to bottom, previous page); SLA map of July 4th, 2007 (day 640 of the mission -this page). The trajectory of float 2900392 for year 2 is superimposed. Dots indicate the position of the float every 5 days. Mission days are indicated next to the dots.

Fig. 11 .

 11 Fig. 11. From top to bottom, sea-level anomaly (SLA) maps of 22 November 2006, 17 January 2007 and 4 July 2007 (days 420, 475 and 640 of the mission). The trajectory of float 2900392 for year 2 is superimposed. Dots indicate the position of the float every 5 days. Mission days are indicated next to the dots.

Figure 12 :Fig. 12 .

 1212 Figure 12: Vertical sections of temperature (in o C, left) and of salinity (right) along float 2900392 trajectory for year 2.

Fig. 13 .

 13 Fig. 13. Trajectory of float 2900104 in the Arabian Sea, superimposed on bathymetry, year 1; the black dot is the starting point of the trajectory; positions every 20 days are indicated by a red dot.

Figure 14 :

 14 Figure 14: Vertical sections of temperature (in o C, left) and of salinity (right) along float 2900104 trajectory for year 1.

Fig. 14 .

 14 Fig. 14. Vertical sections of temperature (in C, left) and of salinity (right) along float 2900104 trajectory for year 1.

Fig. 15 .

 15 Fig. 15. Altimetric maps of 6 September 2006 (day 189 of the mission) and of 7 January 2007 (day 304 of the mission), with the trajectory of float 2900104 superimposed. Dots indicate the position of the float every 10 days.

Figure 16 :Fig. 16 .

 1616 Figure 16: Trajectory of float 2900104 in the Arabian Sea, superimposed on bathymetry, year 2. A red dot indicates the position of the float every 20 days.

Figure 17 :

 17 Figure 17: Vertical sections of temperature (in o C, left) and of salinity (right) along float 2900104 trajectory for year 2.

Float

  1900438 was launched November 18th, 2008 at 56 o 20 E, 16 o 04 N with a 4-day cycle, and its validated data ended March 14th, 2009, at 54 o 26 E, 16 o 02 N. After that date, a 5 month long data gap occurs corresponding to a shift in position by 7 degrees east.

Fig. 17 .

 17 Fig. 17. Vertical sections of temperature (in C, left) and of salinity (right) along float 2900104 trajectory for year 2.

Figure 18 :Fig. 18 .

 1818 Figure 18: Trajectory of float 1900438 in the Arabian Sea, superimposed on bathymetry. A green dot indicates the initial position of the float; red dots mark the position of the float every 20 days.

Figure 20 :Fig. 19 .

 2019 Figure 19: Angle between the isobaths and the trajectory of float 1900438 with respect to time

Figure 19 :Figure 20 :Fig. 20 .

 192020 Figure 19: Angle between the isobaths and the trajectory of float 1900438 with respect to time

Fig. 21 .

 21 Fig. 21. From top to bottom, sea-level anomaly (SLA) maps of 24 December 2008, 28 January and 19 February 2009 (days 36, 72 and 92 of the mission) with the trajectory of float 1900438 superimposed. Dots correspond to the position of this float every 4 days. Mission days are indicated next to the dots.

Figure 22 :Fig. 22 .

 2222 Figure 22: Trajectory of float 2900394, for year 1 from day 50, superimposed on bathymetry. A green dot indicates the initial position of the float; red dots mark the position of the float every 50 days.

Figure 23 :

 23 Figure 23: Vertical sections of temperature (in o C, left) and of salinity (right) along float 2900394 trajectory for year 1 from day 50.

Fig. 23 .

 23 Fig. 23. Vertical sections of temperature (in C, left) and of salinity (right) along float 2900394 trajectory for year 1 from day 50.

Fig. 24 .

 24 Fig. 24. Trajectory of float 2900394, for year 2, superimposed on bathymetry. A green dot indicates the initial position of the float; red dots mark the position of the float every 50 days.

Figure 25 :

 25 Figure 25: Vertical sections of temperature (in o C, left) and of salinity (right) along float 2900394 trajectory for year 2.

Fig. 25 .

 25 Fig. 25. Vertical sections of temperature (in C, left) and of salinity (right) along float 2900394 trajectory for year 2.

Figure 3 . 4 :

 34 Figure 3.4: Maps of the climatological salinity and temperature from GDEM modified by ARGO floats at 600 m depth during the winter monsoon in the western Arabian Sea (up, left and right)); and associated representivity via the ARGO floats (bottom, left); MADT anomaly from AVISO over the spring intermonsoon from 1992 to 2014.

Figure 3 . 5 :

 35 Figure 3.5: Maps of the climatological salinity and temperature from GDEM modified by ARGO floats at 600 m depth during the spring intermonsoon (up, left and right)); and associated representivity via the ARGO floats (bottom, left); MADT anomaly from AVISO over the spring intermonsoon from 1992 to 2014.

4. 5 .Fig. 1

 51 Fig. 1 Map of the Sea of Oman and Northern Arabian Sea bathymetry; the green box is the domain of evolution of float 2900804, the red box that of float 2901370, and the blue box that of float 6900902

3. 2

 2 Analysis of Argo float 2901370 Float 2901370 started its mission November 22, 2009 near 23.78 • N, 60.80 • E, and the last data were collected for this study, June 3, 2012, when the float was near 24.51 • N, 58.02 • E, with a parking depth at 1000 m. The whole float trajectory displayed many loops which ranged in size from 25 to 200 km. The loops in the western Sea of Oman are shown in detail below.

Fig. 2

 2 Fig. 2 Trajectory of Argo float 2900804 superimposed on MADT anomaly maps time-averaged over each period; (top) period of March 3-April 16, 2008; (bottom) period of May 22-June 23, 2008, very high salinities (above 37) were measured along this part of the trajectory

4. 5 .

 5 Selected trajectories and hydrological sections of ARGO floats in the Sea of Oman

Fig. 3

 3 Fig. 3 Time-versus-depth diagram of salinity, as measured by Argo float 2900804, along its trajectory in the vicinity of Ras al Hamra; zoom on the March event (top); zoom on the May-June event (bottom)

3. 3 Fig. 4 5 .Fig

 345 Fig. 4 Trajectory of Argo float 2901370 superimposed on bathymetry; from April 2 to June 30, 2011 (this page, top); from June 30 to December 14, 2011, salinities above 36.5 were measured all along this trajectory(this page, bottom); from May 2 to June 11, 2012 (next page)

Fig. 5 5 .FigFig 5 .

 555 Fig. 5 Trajectory of Argo float 2901370 superimposed on MADT anomaly maps near Ras al Hamra; top) period of May 16-June 17, 2011; (bottom) period of April 4-24, 2012

  in a more transient manner in 1999, 2001, and 2002. However, the density of Argo floats was sufficient to detect PGW ejection at Ras al Hamra only after 2007. Thus, we use altimetry and hydrology of spring 2008, 2011, and 2012 to show how the dipolar eddy velocity extracted PGW from the slope currents. Furthermore, since the dynamical patterns were similar in 2008, 2011, and 2012, we present figures only for March 2008.

  Fig. 10 Horizontal map of salinity in the Sea of Oman at 250-m depths, for spring: from climatology (this page, top); from climatology modified by Argo float data over the 2002-2012 period (this page, bottom); from climatology modified by Argo float data for years 2008, 2011, and 2012 (next page). Each cross represents an individual float profile

Fig

  Fig. 12 (continued)

d

  with β the meridional gradient of the Coriolis parameter, k the wave number magnitude (here the zonal wave number), and R d the Rossby deformation radius. Estimates for these parameters are β = 2 10 -11 m -1 s -1 , k = 2π/(500 km), and the first internal radius of deformation is R d = 40 km.

5Fig. 13 FigFig

 13 Figure 15 displays the wind stress curl for December 2011 and March 2012 (the high-resolution wind data have been available only since 2010). In December, this curl is positive near the Iranian coast and negative near Ras al Hadd. In March, the curl is negative near Ras al Hamra. During the other months, the curl is essentially positive in the Sea of Oman.

Fig

  Fig.16 (continued)

Figures

  Figures 18 and 19 show the evolution of ADT for each month, time-averaged over 3 years: 2008, 2011, and 2012.

Fig

  Fig. 17 Error map for sea surface elevation, on March 12, 2008, in the Sea of Oman

Fig. 18

 18 Fig. 18 Time average of ADT over January (top, left); February (top, right); March (center, left); April (center, right); May (bottom, left); June (bottom, right)

Fig. 19

 19 Fig. 19 Time average of ADT over July (top, left); August (top, right); September (center, left); October (center, right); November (bottom, left); December (bottom, right)

Figure 4 . 1 :

 41 Figure 4.1: Map on the three first EOFs from AVISO altimetry over March 1992 to 2014 with their associated representativity for each year.

  Chapter4. A seasonal dipolar eddy near Ra's Al Hamra (Sea of Oman)

Figure 4 . 2 :

 42 Figure 4.2: Climatological maps in the Sea of Oman for March : salinity and temperature from GDEM modified by ARGO floats at 200 m depth during March (up, left and right)); and associated representativity via the ARGO floats (bottom, left); averaged sea surface temperature from 2002 to 2011 via AVHRR-AMSR measurements (bottom, right).

Figure 1 .

 1 Figure 1. Maps of the northwestern Indian Ocean. Top: area covered by the HYCOM model, including the Arabian Sea and the adjacent gulfs. Bottom: areas of interest in this paper, the Arabian Sea, the Sea of Oman and the main geographical locations.

Figure 2 .

 2 Figure 2. Bi-monthly averaged MADT anomaly (see data and method section for computation) using AVISO data, over years 1992 to 2014. Positive values of the anomaly correspond to an anticyclonic motion, whereas negative ones correspond to a cyclonic motion. Top: January and February (left), March and April (right); center: May and June (left), July and August (right); bottom: September and October (left), November and December (right).

Figure 3 .

 3 Figure 3. Spectra of 22 years of sea-surface height (SSH) (red, (m) 2 ) and 40 years of wind stress (blue, (N m 2 ) 2 ). The spectra are averaged over 15-27 N, 54-75 E. The periods for the major spectral peaks are indicated.

Figure 4 .

 4 Figure 4. Temperature versus salinity diagram in the Sea of Oman (left) and in the northern Arabian Sea (right), from the ARGO float (top) and the HYCOM simulation (bottom). The different water masses are Arabian Sea High Salinity Water (ASHSW), Persian Gulf Water (PGW), Indian Ocean Central Water (IOCW), Indian Ocean Deep Water (IODW), and Antarctic Bottom Water (AABW).

Figure 5 .

 5 Figure 5. Maps from HYCOM during the winter monsoon, showing the presence of the coastal current: (top left) MADT anomaly; (top right) surface velocity; (bottom) anomaly of Ertel potential vorticity. The anomalies of altimetry and Ertel potential vorticity have an opposite sign along the coast, in the current, and offshore. The position of the cross section is indicated on the upper left map.

Figure 6 .

 6 Figure 6. Section across the coastal current during the winter monsoon (from HYCOM): (left) velocities across the section (positive toward the northeast); (right) anomaly of Ertel potential vorticity. A vertical influence of the surface structures is observable. An inversion of the Ertel potential vorticity sign occurs between the surface and depth.

Figure 7 .

 7 Figure7. Maps of the MADT anomaly from AVISO, during the winter monsoon, in January 2011, every 5 days. The first map shows the presence of meanders on the coastal current, then forming coherent mesoscale eddies at 18 N, 58 E on 20 January and at 18 N, 58 E on 30 January.

Figure 8 .Figure 9 .

 89 Figure 8. Maps from HYCOM showing the presence of mesoscale eddies: (upper left) MADT anomaly; (upper right) surface velocity; (lower panel) anomaly of Ertel potential vorticity. The anomalies of Ertel potential vorticity have an opposite sign inside and at the periphery of the eddies. This is primarily seen around the anticyclone north of Ra's Al Hadd, the cyclone interacting with the coastal negative anomaly of Ertel potential vorticity.

Figure 10 .

 10 Figure 10. Maps from HYCOM: (top) correlation coefficient between surface and 600 m depth velocities; (bottom, left) zoom on a mesoscale eddy; (bottom, right) attenuation coefficient between the surface and 600 m depth velocities: up to 50 % of the intensity is observed.

Figure 11 .

 11 Figure 11. Hovmoller diagram in longitude vs. time for the MADT anomaly from AVISO at 16 N (left) and 20 N (right).

Figure 12 .

 12 Figure 12. Correlation between the MADT from AVISO and the surface elevation anomaly created by the wind stress (top) and by the waves along the eastern coast (bottom) using the Fu and Qiu (2002) method.

Figure 13 .Figure 14 .Figure 15 .

 131415 Figure 13. Evolution of the relative vorticity of a Rossby wave over a seamount, located at the middle of the map.

4 )Figure 16 .

 416 Figure 16. Comparison of the tendencies of transfer terms (energy and wind work) with the tendency of the eddies' depth. Five different stationary eddies were studied in the HYCOM simulation. The parameters are the wind work P down (bottom, left), the transfer term from the baroclinic energy to barotropic due to vorticity K 0 K ⇠ (bottom, right), and the sum of those two (top).

  www.ocean-sci.net/11/667/2015/ Ocean Sci., 11, 667-693, 2015 5.7. Evolution of the Persian Gulf Water in the Sea of Oman

Figure 17 .Figure 18 .

 1718 Figure 17. Evolution of the maximal PGW salinity (left) and MADT anomaly (right) during a pulse of PGW and an ejection of PGW at Ra's Al Hadd in year 9 of the HYCOM simulation, on 25 July (first row), 3 August (second row), 12 August (third row) and 21 August (fourth row).

Figure 19 .

 19 Figure 19. Evolution of the maximal PGW salinity (left) and MADT anomaly (right) during the formation of a lee eddy in year 11 of the HYCOM simulation, on 11 November (first row), 19 November (second row), 29 November (third row) and 10 December (fourth row).

Figure 20 .

 20 Figure20. Evolution of the PGW characteristics during the formation of a lee eddy (on 11 November, left) and after its ejection from the coast (on 29 November, right). The first row shows the pathway of the maximal salinity of PGW and the width of the outflow. The second row displays the depth and thickness of the outflow and the third row maximal salinity and transport of salt along the downstream bathymetry.

  3.Between 200 and 400 km, the coastal outflow keeps providing saline water, with a positive transport of salt around 10 8 kg s 1 . www.ocean-sci.net/11/667/2015/ Ocean Sci., 11, 667-693, 2015 5.7. Evolution of the Persian Gulf Water in the Sea of Oman

Figure 21 .

 21 Figure 21. Evolution of the maximal PGW salinity (left) and MADT anomaly (right) during an ejection of PGW off Ra's Al Hamra in year 9 of the HYCOM simulation, on 11 May (first row), 23 May (second row), 4 June (third row) and 12 June (fourth row).

Figure 22 .

 22 Figure 22. Evolution of the PGW characteristics during the ejection off Ra's Al Hamra (on 23 May, left) and after (on 4 June, right). The first row shows the pathway of the maximal salinity of PGW and the width of the outflow. The second row displays the depth and thickness of the outflow and the third row maximal salinity and transport of salt along the downstream.

Figure 23 .

 23 Figure 23. Evolution of the maximal PGW salinity (left) and MADT anomaly (right) during an ejection along the northern coast of the Sea of Oman in year 9 of the HYCOM simulation, on 29 March (first row), 8 April (second row), 14 April (third row) and 30 April (fourth row).

Figure 24 .

 24 Figure24. Evolution of the PGW characteristics during the ejection from the northern coast (on 8 April, left) and after (on 30 April, right). The first row shows the pathway of the maximal salinity of PGW and the width of the outflow. The second row displays the depth and thickness of the outflow and the third row maximal salinity and transport of salt along the downstream.

Figure 5 . 2 :

 52 Figure 5.2: First three EOFs for the winter monsoon (left) and spring inter monsoon (right).

  for the research . . . . . . . . . . . . . . . . . . . . . 6.2 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.3 Data and methods . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.3.1 Climatologies and long time observations . . . . . . . . . . . . . 6.3.2 The Phys-Indien 2011 measurements . . . . . . . . . . . . . . . . 6.3.3 Thermodynamical and dynamical quantities derived from the measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.4 Mesoscale situation from late 2010 to mid 2011 . . . . . . . . . 6.5 Vertical structure of mesoscale eddies in the region in Sping 2011, and relation to PGW . . . . . . . . . . . . . . . . . . . . . 6.5.1 Arabian Sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.5.2 Sea of Oman . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.6 PGW characteristics, pathway and submesoscale structures . . 6.6.1 Submesoscale lens o↵ Ra's Al Hadd . . . . . . . . . . . . . . . . 6.6.2 Recurrence of PGW lenses in the region . . . . . . . . . . . . . . 6.7 Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.8 References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

  and L'Hégaret et al. (2013)) and HYCOM numerical simulations (see L'Hégaret et al. (2015)) reveal different offshore ejection mechanisms of PGW under the influence of the mesoscale eddies, in the Sea of Oman. This occurs particularly 6.2. Introduction in spring when a dipole advects this water mass between the Strait of Hormuz and Ra's al Hamra and along the coast of Iran. In spring 2011, the PhysIndien experiment was carried out around the Arabian Peninsula, recording the thermohaline and dynamical characteristics of the upper ocean. Submesoscale fragments (and in particular a lens) of Persian Gulf Water were sampled in the Sea of Oman and off Ra's Al Hadd, in the Arabian Sea.

5

  Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring

  tendency with two peaks, one in summer and a second one in winter, with the lowest values observed in late September 2010 and from February to mid April 2011. The variations of the wind stress curl displayed the winter plateau from November 2010 to the end of January 2011 and the strongest values starting from late May. Both graphs indicate a spring inter monsoon taking place from February to May 2011. Furthermore, the winds stress curl off Ra's Al Hadd (see figure 4, 10 6.4. Mesoscale situation from late 2010 to mid 2011 right) changed sign during this period, being positive from mid February to mid March, and then negative until May. A positive slope in the wind work, linked to a negative slope of wind stress curl, then deepens the vorticity distribution for surface eddies (see L'Hégaret et al. (2015) and Vic et al. (2014)). On figure 5, the MADT anomaly is displayed over the Sea of Oman and the northwestern Arabian Sea from the winter monsoon to the beginning of the summer monsoon (November 2010 to June 2011). In November 2010, an alongshore current, associated with a positive MADT anomaly, flowed along the southern coast of Oman (L'Hégaret et al. (2015)); it was driven by the Ekman currents. This current started to meander, and the wavelengths were about 2pR d where R d is the first baroclinic radius of deformation (40 km). Offshore of this coastal current, mesoscale cyclones (C1) exitted the Sea of Oman and propagated southwestward along the coast of Oman. Their surface temperature is colder than that of the surrounding waters. The Owen fracture zone and the coastal current can jointly channel these cyclones along the coast (C1 and C2). From December 2010 through February 2011, the MADT anomaly was positive in the Sea of Oman, with the onset of a large anticyclonic eddy (A1), and higher surface temperature whereas colder waters exitted through the Strait of Hormuz. This anticyclone was part of a dipole, located nearly every year in spring near Ra's al Hamra (see L'Hégaret et al. (2013)). South of Ra's Al Hadd, in the Arabian Sea, the alongshore current formed an anticyclone (A2), which splitted apart the C1 and C2 cyclones at Ra's al Hadd. The first cyclone (C1) remained east of Ra's Al Hadd until April; the second part of the cyclone (C2) drifted southwestward, slowly warming up and 11 Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring decreasing in intensity. From March 2011 through May 2011, A1 splitted in two anticylones, A1 and A3. Three main vortices line up along 61 o E, south and north of Ra's Al Hadd: the anticyclone at the mouth of the Sea of Oman (A3), the cyclone at Ra's al Hadd (C1) and the anticyclone south of it (A2).

4 .

 4 Vertical structure of mesoscale eddies in the region in Spring 2011, and relation to PGW This section focus on the vertical characteristics of the eddies presented earlier and their relations to the PGW, through the Physindien experiment results. Figure 6, presents the surface fields of interest during the period of measurements, from the 16 to 30 March 2011; from the MADT, the relative vorticity, Okubo-Weiss parameter and geostrophic velocities are calculated through 12 6.5. Vertical structure of mesoscale eddies in the region in Sping 2011, and relation to PGW derivations. The positions of three sections of interest are indicated on these maps. A first one focuses on the Arabian Sea (AS section), a second one on the Sea of Oman (SO section) and a last one crosses a submesoscale off Ra's Al Hadd (lens section). These sections present velocities measured with the VM-ADCP and geostrophic velocities, density and spice, Okubo-Weiss and Ertel Potential Vorticity fields (see section 2 for computation).

  , bottom) show three water masses. The upper layer, down to about 100 m depth, contains Arabian High Salinity Water (ASHSW), with temperature between 22 and 26 o C, salinity above 36.5 psu and potential density below 25 kg.m 3 . ASHSW is the saltiest and spiciest water recorded along this section. Persian Gulf Water (PGW) is found between about 200 and 350 m with temperature ranging from 17 to 20 o C and salinity varying between 36.2 and 36.4 psu, with a potential density between 25.9 and 27 kg.m 3 . The third water mass is the Indian Ocean Central Water (IOCW), observed between the two salty water masses and below the PGW; with temperature between 14 and 18 o C and salinity between 35.6 and 36.1 psu.

  around anticyclone A1 from 100 m to 350 m depths, colocalized with maximal positive values of the Okubo-Weiss field. Maximal deformation created by A2 tore and wrapped IOCW around the eddy. An other layer of cold IOCW was recorded around C2, between 16.5 o N and 17 o N.

  figure 9, top), but a secondary maximum was found west of 57.5 o E; intense velocities near the

  water patches have thermohaline values characteristic of mixing with the oceanic water, thus indicating a change of direction of the flow. This change occurred offshore of Ra's Al Hamra, as 18 6.5. Vertical structure of mesoscale eddies in the region in Sping 2011, and relation to PGW described in a previous section.

  figure13), looping cyclonically from March to May, then moving northwestward, towards the

  b. Recurrence of PGW lenses in the regionNumerous vertical sections from the ARGO floats(WHOI numbers 1901187, 1901202 and 6900902) revealed the presence of PGW at the mouth of the Sea of Oman. Until the early summer monsoon, localized patches of water with salinity above 37.2 psu, temperature around 20 o C and s 0 ⇡ 26.5 kg.m 3 are observed, with a spacing of the isopycnals above and below. These lenses are found either off Ra's Al Hadd or off the Sea of Oman. In June 2011, anticyclone A3 was advected northward as the summer monsoon began, thus reducing the deformation field off Ra's Al Hamra. Since a strong shear and strain is necessary to break the PGW outflow and to form lens and filaments, it is logical to observe fewer submesoscale PGW structures during the summer monsoon.25Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring6. ConclusionThe Phys-Indien experiment took place in March 2011, during the spring inter monsoon. The surface signature was dominated by mesoscale eddies along the western coast of the Arabian Sea. A strong mesoscale anticyclone was observed in the eastern part of the Sea of Oman, characteristic of the spring inter monsoon; it was associated with a coastal ejection of PGW off Ra's Al Hamra.

Fig. 1 .Fig. 2 .Fig. 3 .Fig. 4 .Fig. 5 .Fig. 6 .Fig. 7 .Fig. 8 .Fig. 10 .Fig. 12 .

 123456781012 Fig. 1. Topographic map of the Arabian Sea and Sea of Oman with the locations of interest. . . . 35 Fig. 2. Maps of the salinity (left) and temperature (center) from GDEM modified by ARGO floats at 250 m depth for a climatological month of March; associated number of floats is indicated on the right-hand panel. . . . . . . . . . . . . . . . . . . . . 36 Fig. 3. Location of the PhysIndien 2011 measurements. The red line follows the VM-ADCP and SeaSoar lines (with the exception of the Red Sea and Gulf of Aden where only the VM-ADCP was activated). The blue crosses and green circles represent the positions of XBT-XCTD casts and of CTDL-ADCP stations respectively. The black crosses represent the launch positions of floats (Surdrift and PROVOR). . . . . . . . . . . . . . 37 Fig. 4. Daily wind stress (left) and wind stress curl (right) in blue and their 15 day means in red, over a 2 o square region off Ra's Al Hadd. . . . . . . . . . . . . . . . 38 Fig. 5. Maps of the MADT (altimetric) anomaly averaged over each month, from November 2010 to June 2011. . . . . . . . . . . . . . . . . . . . . . . 39 Fig. 6. Maps of the surface fields averaged from 16 to 30 March 2011, during the measurements south of Ra's Al Hadd (AS section), in the lens off Ra's Al Hadd (lens section), and across the Sea of Oman (SO section). Fields are : MADT anomaly (top, left); relative vorticity (top, right); Okubo-Weiss criterion (bottom, left); geostrophic velocity (bottom, right). . . . 40 Fig. 7. AS section, south of Ra's Al Hadd from surface down to 350 m depth. Measurements are : VM-ADCP and geostrophic velocities (from SeaSoar density), positive towards the west; Okubo-Weiss paramater calculated from the ARGO-GDEM density and velocity fields and Ertel Potential Vorticity from VM-ADCP and SeaSoar fields; s 0 potential density and spiciness. . . . . . . . . . . . . . . . . . . . . . . . 41 Fig. 8. Potential temperature over salinity profiles (right) in the western Arabian Sea at various locations of interest (left). From north to south : -Yellow diamond : Salty injection around the anticyclone A2; -Cyan diamond : Fresh injection around the anticyclone A2; -Red circle : inside the anticyclone A2; -Blue circle : inside the cyclone C1; -Green cross, at the periphery of the anticyclone; -Red circle : inside the lens off Ra's Al Hadd. . . . . . . 42 Fig. 9. SO sectionin the Sea of Oman from surface down to 350 m depth. Measurements are : VM-ADCP and geostrophic velocities (from SeaSoar density), positive towards the north; Okubo-Weiss paramater calculated from the ARGO-GDEM density and velocity fields and Ertel Potential Vorticity from VM-ADCP and SeaSoar fields; s 0 potential density and spiciness. . . . . . . . . . . . . . . . . . . . . . . . . . 43 Fig. 10. Potential temperature over salinity profiles (right) and their locations of interest (left); in the western Sea of Oman (up) and eastern Sea of Oman (down). Top panel : -Blue diamond : Profile out of the salty outflow; -Green cross : 58 o E lens; -Yellow circle : southern profile in the PGW outflow along the coastal slope; -Grey/Black cross : 57.5 o E lens; -Red circle : northern profile in the PGW outflow along the coastal slope; -Brown circle : cascading PGW outflow. Bottom panel : -Yellow circle : profile inside the anticyclone A3; -Blue diamond : Low spice profile; -Green diamond : High spice profile; -Cyan cross : at the periphery of the anticyclones A1 and A3; -Red circle : profile inside the anticyclone A1. . . 44

FIG. 6 :FIG. 7 :

 67 FIG. 6: Maps of the surface fields averaged from 16 to 30 March 2011, during the measurements south of Ra's Al Hadd (AS section), in the lens off Ra's Al Hadd (lens section), and across the Sea of Oman (SO section). Fields are : MADT anomaly (top, left); relative vorticity (top, right); Okubo-Weiss criterion (bottom, left); geostrophic velocity (bottom, right).

FIG. 9 :

 9 FIG. 9: SO sectionin the Sea of Oman from surface down to 350 m depth. Measurements are : VM-ADCP and geostrophic velocities (from SeaSoar density), positive towards the north; Okubo-Weiss paramater calculated from the ARGO-GDEM density and velocity fields and Ertel Potential Vorticity from VM-ADCP and SeaSoar fields; s 0 potential density and spiciness.

  3.1. Presentation of Dipoles and Summary of Their InteractionsFirst of all, we present here the known positions and evolutions of meddies and cyclones at three different periods: in November 2000,February and June 2001, during Semane 2000, MedTop and Semane 2001,

Figure 1 .

 1 Figure 1. Hydrological sections of the MedTop and Semane experiments superimposed on bathymetry; XBT, XCTD, and CTD casts are indicated in green, red, and yellow.

Figure 2 .

 2 Figure 2. Sketches of the positions of the eddies at different times (a) in November 2000 with M1-C1 and RF1 drift (b) in February 2001 with M2-C2 and M1 relative position, and RF2, RF3 drifts, and (c) in June 2001 with M3-C3 with the previous drifts (RF1, RF2, and RF3).

3. 2 .Figure 3 .

 23 Figure 3. Temperature and salinity section of the meddy M1 of Semane 2000 (November 2000) at its initial position near 35 N and 11 W [from[START_REF] Carton | Meddy dynamics and interaction with neighboring eddies southwest of Portugal: Observations and modeling[END_REF]. The distance is measured from the beginning of the section at its southwesternmost point. This anticyclone has two cores, one with a temperature maximum 13.4 C near 700 m depth, and one with a salinity maximum of 36.8 near 1300 m.

Figure 4 .

 4 Figure 4. Trajectories of Rafos float RF1 of Semane 2000 program starting on November 2000 on the northwest, colored by temperature. Temperature and distance to the meddy center (M1) of RF1 vary very little during the 9 months of record.

Figure 5 .Figure 6 .Figure 7 . 5 .

 5675 Figure 5. (a) Temperature and salinity distributions in (left) meddy M2 (between 5 and 45 km) and (right) cyclone C2 (from 60 km to the end of the section) on the transverse MedTop in February 2001 CTD cross section and (b) across the MW undercurrents on the (right) slope and (left) cyclone C2 (between 5 and 45 km). The distance is measured from the beginning of the section (red square). The anticyclone M2 has two cores, a first with a temperature maximum of 13.4 C near 900 m depth and a second with a salinity maximum of 36.8 psu between 1000 and 1400 m depths. The cyclone C2 has thermohaline maxima of 11.4 C and 36.4 psu near 1200 m depth.

Figure 8 .

 8 Figure 8. Meridional and zonal sections of temperature and salinity in the southern Gulf of Cadiz, during Semane 2001, in June 2001. (a) Quasi-meridional section across the meddy M3, between 0 and 35 km and cyclone C3, between 50 and 110 km; (b) zonal section of the meddy M3 at 34 20 0 N; (c) zonal section of the cyclone C3 at 35 N. The anticyclone has a salinity maximum of 36.6 psu between 1000 and 1500 m depths, the temperature maximum lies between 600 and 800 m depth with 13.5 C.

  Figure 8. (continued)

Figure 9 .

 9 Figure 9. Horizontal maps of (left) temperature and (right) salinity in the southern Gulf of Cadiz, during Semane 2001, in June 2001. (top) At r 1 5 31.9 kg m 23 , the upper core of M3 (south) and the core of C3 (north) show strong thermohaline anomalies, separated by a strip of fresher and colder water. (middle) On isopycnal r 1 5 kg m 23 , out of the eddy cores, temperature, and salinity anomalies decrease. (bottom) On r 1 5 32.2 kg m 23 , higher temperature is observed in M3 than in C3, due to turbulent diffusion of NACW and MW.

5 .

 5 Analysis of the experiments data showing the eddy interaction 3.3. Detailed Analysis of the Eddy Interaction, From Float Trajectories and Altimetric Data

Figure 10 .Figure 11 .

 1011 Figure10. Superposition of trajectories of Rafos floats RF1 (green), RF2 (black), and RF3 (red, except blue on the last map) on altimetric maps (Sea Level Anomaly in cm); from left to right and from top to bottom, months are March to August 2001. For July and August, the green trajectory corresponds to buoy SF.

Figure 12 .Figure 13 .

 1213 Figure 12. Trajectory of a MARVOR (multicycle acoustic float) drifting at 450 dbar, from end of June until end of September 2001 superimposed on SLA. The float rotates around a strong negative altimetric anomaly all along its trajectory while spinning in anticyclonic loops. It suggests that the float is following a meddy rotating around a strong cyclone.

Figure 14 .

 14 Figure14. Ratio of the floats initially in M2 and finally found in M1 (''M1gainedM2''), and of floats initially in M1 lost to the surrounding flow (''M1lost'') with respect to the distances D f to the meddy center.

  20

  ). The Phys-Indien 2011 Experiment . . . . . . . . . . . . . . . . . 31 2.1.1 SeaSoar and Probes . . . . . . . . . . . . . . . . . . . . . . . . . 31 2.1.2 VM-ADCP . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34 2.1.3 Floats . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 2.2 Climatologies and long time observations . . . . . . . . . . . . . 38 2.2.1 ARGO floats and Climatology . . . . . . . . . . . . . . . . . . . 38 2.2.2 Altimetry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38 2.2.3 Atmospheric forcing . . . . . . . . . . . . . . . . . . . . . . . . . 43 2.3 Numerical modelling . . . . . . . . . . . . . . . . . . . . . . . . . 43

	Chapter 2
	Data collection and Processing
	Contents
	2.1

  Also, high frequency winds data are obtained from the European Center for Medium-Range Weather Forecast model (hereafter ECMWF) from the ERA-Interim reanalysis of both model and integrated observations, with a 0.7 ⇥ 0.7 degree resolution and a 12 hours recursion. This resolution is low but su cient enough to resolve the mesoscale and, moreover, provided a long range of data from 1979 to 2014, thus allowing to compute climatological maps of the wind intensities, wind stress curl and Ekman currents. Net surface heat fluxes data are also retrieved from the ECMWF database. For finer scale, but for a shorter period of time, the ASCAT data, obtained from the Centre de recherche et d'Exploitation Satellitaire (CERSAT) at Ifremer, with a 1/4 ⇥ 1/4 degrees resolution and daily field from 2007 to 2014.The bathymetry ETOPO1 map is obtained from the NGDC data center with a very high resolution of 1/60 o . Data can be retrieved on http://www.ngdc.noaa.gov/mgg/global/

	2.3. Numerical modelling
	2.2.3 Atmospheric forcing

  Motivations for the research . . . . . . . . . . . . . . . . . . . . . 3.2 Abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.3 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.4 Data and methods . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.5 Gulf of Aden . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Gulf of Aden . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69 3.11.2 Arabian Sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75

	Chapter 3
	Mesoscale variability of water
	masses in the Arabian Sea as
	revealed by Argo floats
	Contents
	3.1 3.5.1 Float 1900432 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49
	3.5.2 Float 2900392 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51
	3.6 Arabian Sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
	3.6.1 Float 2900104 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54
	3.6.2 Float 1900438 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57
	3.6.3 Float 2900394 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57
	3.7 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
	3.8 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
	3.9 Appendix A : Climatological maps . . . . . . . . . . . . . . . . .
	3.10 References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
	3.11 Further discussion . . . . . . . . . . . . . . . . . . . . . . . . . . .
	3.11.1

. A total of Ocean Sci., 8, 227-248, 2012 www.ocean-sci.net/8/227/2012/

  Chapter3. Mesoscale variability of water masses in the Arabian Sea as revealed by Argo floats

www.ocean-sci.net/8/227/2012/ Ocean Sci., 8, 227-248, 2012

and one Ocean Sci., 8, 227-248, 2012 www.ocean-sci.net/8/227/2012/

  On day 40 of the mission (16 May), two large gyres, one cyclonic east of 46 30 0 E

	Chapter3. Mesoscale variability of water masses in the Arabian Sea as revealed by Argo
					floats
	Maps of Sea Level Anomaly Merged (cm) on 16 May 2007 (day 40)
					12
	14 o N				10
					8
	13 o N			48
					6
				40
	12 o N		56	32	4
					2
	11 o N
					0
	45 o E	46 o E	47 o E	48 o E
	Maps of Sea Level Anomaly Merged (cm) on 27 Jun 2007 (day 80)
					15
	14 o N				10
	13 o N	80	96		5
			72
			88		0
	12 o N
					-5
	11 o N
					-10
	45 o E	46 o E	47 o E	48 o E
				11
	14 o N
	13 o N
	12 o N
	11 o N
	45 o E	46 o E	47 o E	48 o E

100 108 116 124 Maps of Sea Level Anomaly Merged (cm) on 25 Jul 2007 (day 108)

0 N for Ocean Sci., 8, 227-248, 2012 www.ocean-sci.net/8/227/2012/

  Chapter3. Mesoscale variability of water masses in the Arabian Sea as revealed by Argo floats

		Maps of Sea Level Anomaly Merged (cm) on 22 Nov 2006 (day 420)
								25
	15 o N							20
								15
	14 o N							10
								5
	13 o N	400						0
		410						-5
	12 o N		430
								-10
		420
								-15
	11 o N
								-20
	46 o E	47 o E	48 o E		49 o E	50 o E	51 o E	52 o E
		Maps of Sea Level Anomaly Merged (cm) on 17 Jan 2007 (day 475)
								25
	15 o N							20
								15
	14 o N				455
								10
								485
	13 o N						465	5
							475	0
	12 o N							-5
								-10
	11 o N
								-15
	46 o E	47 o E	48 o E		49 o E	50 o E	51 o E	52 o E
		Maps of Sea Level Anomaly Merged (cm) on 4 Jul 2007 (day 640)
					23
	15 o N
	14 o N
			630	640
	13 o N		620			650
	12 o N
	11 o N
	46 o E	47 o E	48 o E	49 o E	50 o E	51 o E	52 o E

Carton et al.: Mesoscale variability in the Arabian Sea

	Maps of Sea Level Anomaly Merged (cm) on 20 Sep 2006 (day 189)
	17 o N				5
					0
	16 o N				-5
					-10
	15 o N		189		-15
	209		169		-20
	14 o N
					-25
	13 o N				-30
					-35
	54 o E	55 o E	56 o E	57 o E	58 o E
	17 o N
					304
			284
	16 o N
				324
	15 o N
	14 o N
	13 o N
	54 o E	55 o E	56 o E	57 o E	58 o E

www.ocean-sci.net/8/227/2012/ Ocean Sci., 8, 227-248, 2012 3.6. Arabian Sea 238 X. Maps of Sea Level Anomaly Merged (cm) on 10 Jan 2007 (day 304)

A Climatological maps of temperature and salinity at 300 and 600 m depth 244 X. Carton et al.: Mesoscale variability in the Arabian Sea

  Chapter3. Mesoscale variability of water masses in the Arabian Sea as revealed by Argo floats

	Climatological temperature map at 300 m depth in January	19		Climatological salinity map at 300 m depth in January
	24 o N 24 o N	Climatological temperature map at 300 m depth in July	17 18 17	24 o N 24 o N	Climatological salinity map at 300 m depth in July		36.6 36.6 36.8 37 36.8
	20 o N 20 o N						16 16	20 o N 20 o N						36.2 36.4 36.2 36.4
	12 o N 16 o N 12 o N 16 o N						13 13 14 15 14 15	12 o N 12 o N 16 o N 16 o N						35.4 35.6 35.6 35.8 36 35.8 36
	8 o N						12	8 o N						35.2 35.4
	8 o N	48 o E	54 o E	60 o E	66 o E	72 o E	12	8 o N	48 o E	54 o E	60 o E	66 o E	72 o E 35.2
		48 o E	54 o E	60 o E	66 o E	72 o E		48 o E	54 o E	60 o E	66 o E	72 o E
	Climatological temperature map at 300 m depth in March	20		Climatological salinity map at 300 m depth in March
	24 o N Climatological temperature map at 300 m depth in September 18 19 24 o N 17	Climatological salinity map at 300 m depth in September 24 o N 24 o N	37 36.6
	20 o N 20 o N						17 16	20 o N 20 o N						36.5 36.2 36.4
	16 o N						16 15	16 o N						36
	16 o N						15 14	16 o N						36 35.8
	12 o N 12 o N						13 14 13	12 o N 12 o N						35.5 35.6
	8 o N 8 o N	48 o E	54 o E	60 o E	66 o E	72 o E	12 12	8 o N 8 o N	48 o E	54 o E	60 o E	66 o E	72 o E 35.2 35.4
		48 o E	54 o E	60 o E	66 o E	72 o E		48 o E	54 o E	60 o E	66 o E	72 o E
		Climatological temperature map at 300 m depth in May			Climatological salinity map at 300 m depth in May
	20 o N 24 o N Climatological temperature map at 300 m depth in November 17 18 19 17 24 o N 18	20 o N 24 o N 24 o N	Climatological salinity map at 300 m depth in November	36.6 37.2 36.4 36.8 37 36.6
	20 o N						16 16	20 o N						36.4 36.2
	8 o N 12 o N 16 o N 12 o N 16 o N						12 13 14 15 13 14 15	8 o N 12 o N 12 o N 16 o N 16 N o						35.2 35.4 35.6 35.4 35.6 35.8 36 36.2 35.8 36
	8 o N	48 o E	54 o E	60 o E	66 o E	72 o E	12	8 o N	48 o E	54 o E	60 o E	66 o E	72 o E 35.2
		48 o E	54 o E	60 o E	66 o E	72 o E		48 o E	54 o E	60 o E	66 o E	72 o E
							46

Fig. A1. (Continued from previous page) Horizontal maps of temperature (in C, left) and salinity (right) at 300 m depth for January, March, May, July, September, November (from top to bottom), from the climatology. Ocean Sci., 8, 227-248, 2012 www.ocean-sci.net/8/227/2012/ Figure 26: (Continued from previous page) Horizontal maps of temperature (in o C, left) and salinity (right) at 300 m depth for January, March, May, July, September, November (from top to bottom), from the climatology 47 Fig. A1. Continued.

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  Chapter3. Mesoscale variability of water masses in the Arabian Sea as revealed by Argo floats

	Climatological temperature map at 600 m depth in January			Climatological salinity map at 600 m depth in January
		Climatological temperature map at 600 m depth in July	17		Climatological salinity map at 600 m depth in July	37
	24 o N 24 o N						16 17 18	24 o N 24 o N					36.6 36.8 36.8 37
	16 o N 20 o N 16 o N 20 o N						14 13 15 14 15 16	16 o N 16 o N 20 o N 20 o N					35.8 36 36 36.2 36.2 36.4 36.4 36.6
	12 o N						12 13	12 o N					35.6 35.8
	12 o N						11 12	12 o N					35.4 35.6
	8 o N 8 o N	48 o E 48 o E	54 o E 54 o E	60 o E 60 o E	66 o E 66 o E	72 o E 72 o E 10 10 11	8 o N 8 o N	48 o E 48 o E	54 o E 54 o E	60 o E 60 o E	66 o E 66 o E	72 o E 72 o E 35.2 35.2 35.4
	Climatological temperature map at 600 m depth in March			Climatological salinity map at 600 m depth in March
	24 o N 24 o N Climatological temperature map at 600 m depth in September 17 18 19 17 18	24 o N 24 o N Climatological salinity map at 600 m depth in September	36.8 36.8 37 37.2 37
	20 o N						16 16	20 o N					36.6 36.6
	20 o N						15 15	20 o N					36.4 36.4
	12 o N 16 o N 12 o N 16 o N						11 12 12 13 13 14 14	12 o N 12 o N 16 o N 16 o N					35.4 35.6 35.6 35.8 35.8 36 36 36.2 36.2
	8 o N 8 o N	48 o E 48 o E	54 o E 54 o E	60 o E 60 o E	66 o E 66 o E	72 o E 72 o E 10 10 11	8 o N 8 o N	48 o E 48 o E	54 o E 54 o E	60 o E 60 o E	66 o E 66 o E	72 o E 72 o E 35.2 35.2 35.4
		Climatological temperature map at 600 m depth in May			Climatological salinity map at 600 m depth in May
	24 o N 24 o N Climatological temperature map at 600 m depth in November 17 16 17 18	24 o N 24 o N	Climatological salinity map at 600 m depth in November	36.8 37 37 37.2
	20 o N 20 o N						14 15 15 16	20 o N 20 o N					36.2 36.4 36.4 36.6 36.6 36.8
	12 o N 16 o N 12 o N 16 o N						11 12 13 12 13 14	12 o N 12 o N 16 o N 16 N o					35.4 35.6 35.6 35.8 35.8 36 36 36.2
	8 o N 8 o N	48 o E 48 o E	54 o E 54 o E	60 o E 60 o E	66 o E 66 o E	72 o E 72 o E 10 10 11	8 o N 8 o N	48 o E 48 o E	54 o E 54 o E	60 o E 60 o E	66 o E 66 o E	72 o E 72 o E 35.2 35.2 35.4
	Fig. A2. (Continued from previous page) Horizontal maps of temperature (in C, left) and salinity (right) at 600 m depth for January, March,
	May, July, September, November (from top to bottom), from the climatology.
	Figure 27: (Continued from previous page) Horizontal maps of temperature (in
							48

Ocean Sci., 8, 227-248, 2012 www.ocean-sci.net/8/227/2012/ o C, left) and salinity (right) at 600 m depth for January, March, May, July, September, November (from top to bottom), from the climatology 49 Fig. A2. Continued. www.ocean-sci.net/8/227/2012/ Ocean Sci., 8, 227-248, 2012 3.9. Appendix A : Climatological maps 67 3.11. Further discussion

  Chapter3. Mesoscale variability of water masses in the Arabian Sea as revealed by Argo floats Maps of the climatological salinity and temperature from GDEM modified by ARGO floats at 600 m depth during the winter monsoon in the Gulf of Aden (up, left and right)); and associated representivity via the ARGO floats (bottom, left); MADT anomaly from AVISO over the winter monsoon from 1992 to 2014. Maps of the climatological salinity and temperature from GDEM modified by ARGO floats at 600 m depth during the summer monsoon in the Gulf of Aden (up, left and right)); and associated representivity via the ARGO floats (bottom, left); MADT anomaly from AVISO over the summer monsoon from 1992 to 2014. Chapter3. Mesoscale variability of water masses in the Arabian Sea as revealed by Argo floats Maps of the climatological salinity and temperature from GDEM modified by ARGO floats at 600 m depth during the spring intermonsoon in the Gulf of Aden (up, left and right)); and associated representivity via the ARGO floats (bottom, left); MADT anomaly from AVISO over the spring intermonsoon from 1992 to 2014.

	42 o E o N o N o N o N 42 o E o N o N o N o N Figure 3.1: 3.11. Further discussion 44 o E 46 o E 48 o E Map of salinity from the Argo floats at 600 m depth during the winter monsoon 50 o E 52 o E 44 o E 46 o E 48 o E 50 o E 52 o E Map of the 298 Argo measurements distribution and their representativity ranges at 600 m depth during the winter monsoon 44 o E 46 o E 48 o E 50 o E 52 o E 54 o E Map of salinity from the Argo floats at 600 m depth during the summer monsoon 42 o E 10 o N 12 o N 14 o N 16 o N Salinity [psu] 35.6 35.8 36 36.2 36.4 36.6 36.8 37 42 o E 44 o E 46 o E 48 o E 50 o E 52 o E 54 o E 10 o N 12 o N 14 o N 16 o N Map of the 252 Argo measurements distribution and their representativity ranges at 600 m depth during the summer monsoon [%] 0 10 20 30 40 50 60 70 80 90 100 Figure 3.2: 42 o E 44 o E 46 o E 48 o E 50 o E 52 o E o N o N o N o N Map of salinity from the Argo floats at 600 m depth during the spring inter monsoon 54 o E 54 o E 54 o E 42 o E 44 o E 46 o E 48 o E 50 o E 52 o E 54 o E o N o N o N o N Map of the 258 Argo measurements distribution and their representativity ranges at 600 m depth during the spring inter monsoon 54 o E 56 o E 58 o E 60 o E 62 o E 64 o E 66 o E 12 o N 14 o N 16 o N 18 o N 20 o N 22 o N Map of salinity from the Argo floats at 200 m depth during the winter monsoon Salinity [psu] 35.4 35.6 35.8 36 36.2 36.4 14 o N 16 o N 18 o N 20 o N Map of the 1696 Argo measurements distribution and their Figure 3.3: 3.11. Further discussion 22 o N representativity ranges at 200 m depth during the winter monsoon	Salinity [psu] [%] 42 o E 10 o N 35.6 35.8 36 36.2 36.4 36.6 36.8 37 0 10 20 30 40 50 60 70 80 90 100 12 o N 14 o N 16 o N Map of temperature from the Argo floats at 600 m depth during the summer monsoon 42 o E 44 o E 46 o E 48 o E 50 o E 52 o E 10 o N 12 o N 14 o N 16 o N Map of temperature from the Argo floats at 600 m depth during the winter monsoon 54 o E 42 o E 44 o E 46 o E 48 o E 50 o E 52 o E 54 o E 10 o N 12 o N 14 o N 16 o N MADT anomaly during the winter monsoon 44 o E 46 o E 48 o E 50 o E 52 o E 54 o E Temperature [ o C] 11.5 12 12.5 13 13.5 14 14.5 15 15.5 16 16.5 17 42 o E 44 o E 46 o E 48 o E 50 o E 52 o E 54 o E 10 o N 12 o N 14 o N 16 o N MADT anomaly during the summer monsoon MADT anomaly [m] -0.25 -0.2 -0.15 -0.1 -0.05 0 0.05 0.1 0.15 0.2 0.25 Salinity [psu] 35.6 35.8 36 36.2 36.4 36.6 36.8 37 42 o E 44 o E 46 o E 48 o E 50 o E 52 o E 54 o E 10 o N 12 o N 14 o N 16 o N Map of temperature from the Argo floats at 600 m depth during the spring inter monsoon [%] 0 10 20 30 40 50 60 70 80 90 100 42 o E 44 o E 46 o E 48 o E 50 o E 52 o E 54 o E 10 o N 12 o N 14 o N 16 o N MADT anomaly during the spring intermonsoon 54 o E 56 o E 58 o E 60 o E 62 o E 64 o E 66 o E 12 o N 14 o N 16 o N 18 o N 20 o N 22 o N Map of temperature from the Argo floats at 200 m depth during the winter monsoon Temperature [ o C] 14.5 15 15.5 16 16.5 17 17.5 18 18.5 19	Temperature [ o C] MADT anomaly [m] -0.25 11.5 12 12.5 13 13.5 14 14.5 15 15.5 16 16.5 17 -0.2 -0.15 -0.1 -0.05 0 0.05 0.1 0.15 0.2 0.25 Temperature [ o C] 11.5 12 12.5 13 13.5 14 14.5 15 15.5 16 16.5 17 MADT anomaly [m] -0.25 -0.2 -0.15 -0.1 -0.05 0 0.05 0.1 0.15 0.2 0.25
	54 o E 12 o N	56 o E	58 o E	60 o E	62 o E	64 o E	66 o E

  salinity from the Argo floats at 200 m depth during the summer monsoon

									Map of temperature from the Argo floats at 200 m depth during the summer monsoon 22 o N	19
							36.4
									18.5
							36.2		20 o N	18
									17.5
							35.8 36	Salinity [psu]	16 o N 18 o N	16.5 17	C] Temperature [ o
									16
							35.6
									14 o N	15.5
									15
							35.4
									54 o E 12 o N	56 o E	58 o E	60 o E	62 o E	64 o E	66 o E	14.5
	54 o E	56 o E	58 o E	60 o E	62 o E	64 o E	66 o

  Chapter4. A seasonal dipolar eddy near Ra's Al Hamra (Sea of Oman)

	Fig. 15 (Top) Wind stress curl in the Gulf of Oman and the Northern Arabian Sea, in	27 o N	Wind Stress Curl for a month starting on 2011/12/01	8 x 10 -5
	December 2011; (bottom) wind
	stress curl in March 2012	26 o N							6
											4
		25	o N
											2
		24	o N
											0	] [N/m 3
		23 o N
											-2
		22	o N							-4
		21 o N							-6
		20	56 o N o E	o 57 E	o 58 E	o 59 E	o 60 E	o 61 E	o 62 E	o E 63	-8
		20	56 o E	o 57 E	o 58 E	o 59 E	o 60 E	o 61 E	o 62 E	63 o E

Ocean Data View Temperature map at 200 m depth with the ODV interpolation, L=50km

	4.8. Conclusions
	Fig. 16 Temperature field at
	200-m depths obtained by using all float profiles of spring	27˚N 27˚N						22
	2002-2012, as individual
	profiles with the ODV software
	(this page, top); temperature field obtained with the	26˚N 26˚N						21
	interpolation method described
	in this paper (this page, bottom). Number of significative points								20
	map for hydrology,	25˚N 25˚N
	interpolated/extrapolated from
	Argo profiles, for spring (next page)								19
		24˚N 24˚N
									18
		23˚N 23˚N
									17
		56˚E 21˚N 22˚N 56˚E 21˚N 22˚N	57˚E 57˚E	58˚E 58˚E	59˚E 59˚E	60˚E 60˚E	61˚E 61˚E	62˚E 62˚E Ocean Data View	15 16
		22	56 o E	o 57 E	o 58 E		o 59 E	o 60 E	o 61 E	62 o E
									99

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  Chapter5. Mesoscale variability in the Arabian Sea from HYCOM model results and observations : impact on the Persian Gulf Water path

				Cross section of the velocity (perpendicular to the section)							Cross section of the anomaly of Ertel potential vorticity
														0.4
		-50												0.3			-50
		-100												0.2			-100
		-150												0.1			-150
	Depth (m)	-200												0	[m.s -1 ]	Depth (m)	-200
		-250												-0.1		-250
		-300												-0.2		-300
		-350												-0.3		-350
		-400	59.4	59.6	59.8	60	60.2	60.4	60.6	60.8	61	61.2	61.4	-0.4		59.4 -400	59.6	59.8	60	60.2	60.4	60.6	60.8	61	61.2	61.4
							Longitude ( o E)														Longitude ( o E)

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  Chapter5. Mesoscale variability in the Arabian Sea from HYCOM model results and observations : impact on the Persian Gulf Water path

			Map of the MADT anomaly						Surface map of velocity
						0.15								0.7
	26 o N					0.1		26 o N						0.6
	25 o N							25 o N						0.5
						0.05
	24 o N					0	[m]	24 o N						0.4	[m.s -1 ]
	23 o N							23 o N						0.3
						-0.05
	22 o N							22 o N						0.2
	21 o N					-0.1		21 o N						0.1
	56 o E	58 o E	60 o E	62 o E	64 o E	66 o E		56 o E	58 o E	60 o E	62 o E	64 o E	66 o E	0
					Surface map of the anomaly of Ertel potential vorticity
				26 o N
				25 o N
				24 o N
				23 o N
				22 o N
				21 o N
				56 o E	58 o E	60 o E	62 o E	64 o E	66 o E

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		5.6. Interactions of the mesoscale structures
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		Correlation between surface and 600 m depth velocities
	26 o N
	24 o N
	22 o N
	20 o N
	18 o N
	16 o N
	56 o E	60 o E	64 o E	68 o E	72 o E

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  Evolution of the Persian Gulf Water in the Sea of Oman

	680	-6 -4 6 x 10 -8 P. L'-1 -2 0 2 4 Wind work + transfer of energy tendency (kg.s -4 )		-0.8 Wind work + Kinetic energy transfer term / Eddy depth -0.6 -0.4 -0.2 0 0.2 0.4 0.6	0.8	1
												Eddy depth tendency (m.s -1 )	x 10 -3
		-1	-0.8	-0.6	-0.4	-0.2	0	0.2	0.4	0.6	0.8	1
												x 10
												123

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	26 o N		PGW maximal salinity on November, 11 th (year 11)			26 o N		MADT anomaly on November, 11 th (year 11)		0.2
								38
																	0.15
	25 o N								25 o N
																	0.1
								37.5
	24 o N								24 o N							0.05
									[PSU]								0	[m]
								37
	23 o N								23 o N							-0.05
	22 o N							36.5	22 o N							-0.1
																	-0.15
	56 o E o 21 N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	36	56 o E o 21 N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	-0.2
	26 o N		PGW maximal salinity on November, 19 th (year 11)
								38
	25 o N
								37.5
	24 o N
								37	[PSU]
	23 o N
	22 o N							36.5
	56 o E o 21 N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	36	21	56 o E	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E

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  Chapter5. Mesoscale variability in the Arabian Sea from HYCOM model results and observations : impact on the Persian Gulf Water path

	PGW pathway during a lee eddy formation associated with the core and edges salinity
	26 o N								26 o
								39
	25 o N							38.5	25 o N
	150km →
								38
	24 o N	300km →						[PSU]	24 o N
			450km →					37.5
						← 750km
	23 o N								23 o N
								37
				600km →
	56 o E 22 o N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	36.5	56 o E 22 o N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E

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  Chapter5. Mesoscale variability in the Arabian Sea from HYCOM model results and observations : impact on the Persian Gulf Water path

		PGW pathway during an ejection off Ra's Al Hamra
		associated with the core and edges salinity
	26 o N
								39
	25 o N							38.5
	150km →
		300km →
	24 o N	450km → 600km →			1050km →			38	[PSU]
			750km →
								37.5
			900km →
	23 o N
								37
	56 o E 22 o N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	36.5	56 o E 22 o N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E

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  .7. Evolution of the Persian Gulf Water in the Sea of Oman

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	26 o N		PGW maximal salinity on March, 29 th (year 8)			26 o N		MADT anomaly on March, 29 th (year 8)		0.2
								38
																0.15
	25 o N								25 o N
																0.1
								37.5
	24 o N								24 o N							0.05
								37	[PSU]							0	[m]
	23 o N								23 o N							-0.05
	22 o N							36.5	22 o N							-0.1
																-0.15
	56 o E o 21 N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	36	56 o E o 21 N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	-0.2
	26 o N		PGW maximal salinity on April, 8 th (year 8)			26 o N		MADT anomaly on April, 8 th (year 8)			0.2
								38
																0.15
	25 o N								25 o N
																0.1
								37.5
	24 o N								24 o N							0.05
								37	[PSU]							0	[m]
	23 o N								23 o N							-0.05
	22 o N							36.5	22 o N							-0.1
																-0.15
	56 o E 21 o N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	36	56 o E 21 o N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	-0.2
	26 o N		PGW maximal salinity on April, 14 th (year 8)			26 o N		MADT anomaly on April, 14 th (year 8)		0.2
								38
																0.15
	25 o N								25 o N
																0.1
								37.5
	24 o N								24 o N							0.05
								37	[PSU]							0	[m]
	23 o N								23 o N							-0.05
	22 o N							36.5	22 o N							-0.1
																-0.15
	56 o E o 21 N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	36	56 o E o 21 N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	-0.2
	26 o N		PGW maximal salinity on April, 30 th (year 8)			26 o N		MADT anomaly on April, 30 th (year 8)
								38
	25 o N								25 o N
								37.5
	24 o N								24 o N
								37	[PSU]
	23 o N								23 o N
	22 o N							36.5	22 o N
	56 o E 21 o N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	36	56 o E 21 o N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E

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		PGW pathway during an ejection along the iranian coast
			associated with the core and edges salinity
	26 o N
								39
	25 o N			← 300km				38.5
	150km →
			450km →					38
	24 o N							[PSU]
								37.5
			600km →
	23 o N								23 o N
								37
	56 o E 22 o N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E	36.5	56 o E 22 o N	57 o E	58 o E	59 o E	60 o E	61 o E	62 o E

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  Once formed, these eddies usually drift southwestward except for two fairly stationary eddies, one south of Ra's Al Hadd (between 61 o E-63 o E and 21 o N-23 o N), and one o↵ Ra's Madrakah (at 59 o E-60 o E 19 o N-20 o N), which become less intense with time.

  Chapter5. Mesoscale variability in the Arabian Sea from HYCOM model results and observations : impact on the Persian Gulf Water path

				5.11. Further studies : Seasonal variability
			First EOF (29%) for June to September 1992 to 2014 and amplitude First EOF (34%) for November to February 1992 to 2014 and amplitude	0.06	0.05				First EOF (21%) for October 1992 to 2014 and amplitude First EOF (24%) for March to May 1992 to 2014 and amplitude	0.08
	26 o N	26 o N																	0.04	26 o N	26 o N		0.06
																		0.04						0.06
																				0.03					0.04
	24 o N	24 o N															0.02	0.02	24 o N	24 o N		0.04	0.02
	22 o N	22 o N																	0.01	22 o N	22 o N		0.02
																		0		0					0	0
	20 o N	20 o N															-0.02	-0.02 -0.01	20 o N	20 o N		-0.02	-0.02
	18 o N	18 o N																	-0.03	18 o N	18 o N		-0.04	-0.04
																		-0.04
	16 o N	16 o N																	-0.04	16 o N	16 o N		-0.06	-0.06
			56 o E	56 o E	60 o E	60 o E	64 o E	64 o E	68 o E	68 o E	72 o E	72 o E		-0.06	-0.05			56 o E	56 o E	60 o E	60 o E	64 o E	64 o E	68 o E	68 o E	72 o E	72 o E	-0.08
	3		3																		3		2
	2		2																		2
																							1
	1		1																		1
	1990 0	1992 0	1995 1994	1996	2000 1998	2000	2005 2002	2004	2010 2006	2008	2015 2010	2012	1992 0	1994 1990 0	1996	1998 1995	2000	2002 2000	2004	2006 2005	2008	2010 2010	2012	2015
			Second EOF (10%) for June to September 1992 to 2014 and amplitude Second EOF (6.7%) for November to February 1992 to 2014 and amplitude							Second EOF (8.6%) for October 1992 to 2014 and amplitude Second EOF (11%) for March to May 1992 to 2014 and amplitude
																									0.08
	26 o N	26 o N															0.06	0.08	26 o N	26 o N		0.06	0.08
																				0.06					0.06
	24 o N	24 o N															0.04	0.04	24 o N	24 o N		0.04	0.04
	22 o N	22 o N															0.02	0.02	22 o N	22 o N		0.02	0.02
																		0		0					0	0
	20 o N	20 o N															-0.02	-0.02	20 o N	20 o N		-0.02	-0.02
	18 o N	18 o N															-0.04	-0.04	18 o N	18 o N		-0.04	-0.04
																				-0.06					-0.06
	16 o N	16 o N															-0.06	-0.08	16 o N	16 o N		-0.06	-0.08
			56 o E	56 o E	60 o E	60 o E	64 o E	64 o E	68 o E	68 o E	72 o E	72 o E							56 o E	56 o E	60 o E	60 o E	64 o E	64 o E	68 o E	68 o E	72 o E	72 o E	-0.08
	2		1																		2	1.5
	0 1		0																		0	0.5 1
	1990 -1	1992 -1	1995 1994	1996	2000 1998	2000	2005 2002	2004	2010 2006	2008	2015 2010	2012	1992 -2	1994 1990 0	1996	1998 1995	2000	2002 2000	2004	2006 2005	2008	2010 2010	2012	2015
			Third EOF (6.6%) for June to September 1992 to 2014 and amplitude Third EOF (5.2%) for November to February 1992 to 2014 and amplitude	0.08					Third EOF (8.2%) for October 1992 to 2014 and amplitude Third EOF (5.5%) for March to May 1992 to 2014 and amplitude	0.08
																									0.08
	26 o N	26 o N															0.06	0.06	26 o N	26 o N		0.06
																									0.06
	24 o N	24 o N															0.04	0.04	24 o N	24 o N		0.04	0.04
																		0.02	0.02					0.02
	22 o N	22 o N															0		0	22 o N	22 o N		0	0 0.02
	20 o N	20 o N															-0.02	-0.02	20 o N	20 o N		-0.02	-0.02
	18 o N	18 o N															-0.04	-0.04	18 o N	18 o N		-0.04	-0.04
																		-0.06						-0.06
	16 o N	16 o N																	-0.06	16 o N	16 o N		-0.06
			56 o E	56 o E	60 o E	60 o E	64 o E	64 o E	68 o E	68 o E	72 o E	72 o E		-0.08				56 o E	56 o E	60 o E	60 o E	64 o E	64 o E	68 o E	68 o E	72 o E	72 o E	-0.08	-0.08
	2		1																		2		2
	0 1		-1 0																		0		0
	1990 -1	1992 -2	1995 1994	1996	2000 1998	2000	2005 2002	2004	2010 2006	2008	2015 2010	2012	1992 -2	1994 1990 -2	1996	1998 1995	2000	2002 2000	2004	2006 2005	2008	2010 2010	2012	2015

Figure 5.1: First three EOFs for the summer monsoon (left) and the fall inter monsoon (right).

  ASHSW is colder than in the Arabian Sea, between 22 o C and 25 o C and a homogeneous salinity of 36.6 psu. Its bottom depth varied between 75 and 150 m depth. Below, the IOCW is observed with the lowest salinity, between 36.1 and 36.3, and temperature varied 15 o C and 21 o C.As in the Arabian Sea, the depth of equilibrium of the water masses depends of the mesoscale eddies, with a shallowing of the isopycnals between 58 o E and 59 o E, where the relative vorticity map is positive, and a deepening east of 59.3 o E, below the anticyclone (see figure9, bottom).Nevertheless, between 57.5 and 57.75 o E and at 58 o E, the isopycnals were stretched vertically between 250 and 350 m depth, at the position of the highest salinity of PGW, indicating an anticyclonic motion. Different submesoscale structures were observed here, filaments and lenses, described in the next section.

	The PGW was found at two locations with two strong differences.
	A first patch of PGW is observed on the western part of the section, with strong thermohaline
	anomalies, temperature between 21 o C and 22 o C, and salinity beyond 36.8 psu, up to 38 psu, with
	depth varying between 150 and 320 m depth. This layer was fragmented into smaller patches
	with the strongest observed thermohaline characteristics presented on the diagram figure 10; at
	57.25 o E (brown circles profile), between 57.5 and 57.75 o E (grey crosses profile) and between
	57.9 and 58.2 o E (green crosses profile). This water mass ceased at 58.8 o E.
	Oppositely, diluted PGW is observed on the eastern section, with temperature from 16 o C to
	18 o C, salinity of 36.4 psu and between 275 and 350 m depth. These characteristics correspond to
	those measured below the cyclone in the Arabian Sea.

  4 to 36.6 psu, and inside the lens off Ra's Al Hadd with a peak at 37.4 psu. These two peaks correspond to Persian Gulf Water from different seasons. The lighter outflow presents higher temperature and lower salinity (see table1). During the summer monsoon and at the beginning of the winter monsoon, the outflowing Persian Gulf Water is warmer, up to 23 o C in November; it is colder (about 21 o C) from February to May. Salinity also varies with a maximum between December and May and with a peak in April and May, around 38.5 psu. In summer and early winter PGW salinity decreases by about 1 psu, with a value around 37.3 in November. Thus the lightest waters outflow between July and August and in November, while the densest are observed in winter and spring. 19Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring Maps figure 11 indicate the maximal salinity in the PGW layer, and the associated temperature, potential density and spiciness; maps of Eulerian transport (figure

  ). Above the lens, a fresh layer of IOCW is found with temperature between 16 and 19 o C and salinity below 36 psu up to 100 m depth. The surface layer is occupied by ASHSW, with temperature between 26 and 21.5 o C and a salinity of 36.6 psu. At the depth of the lens, and surrounding it, mixed PGW is observed with temperature between 14 and 16 o C and salinity peaking at 36.3 psu. The total salt and of heat content, inside the 36.6 psu (or 18 o C) contour, is 2.59 10 12 kg and 6.54 10 19 J, and the lens volumic transport (across the section) is above 0.4 Sv. The density section on figure14 (bottom,

	Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring
	left) indicates the presence of this anticyclonic lens, with an increased spacing of the isopycnals
	between 250 and 350 m depth. The spiciness presents injections of neighbouring water beneath
	the lens, which are composed of upper IOCW.
	23

  submesoscale lens recorded off Ras Al Hadd possessed strong salinity (over 37 psu) and temperature, characteristic of the winter monsoon. Different hypotheses are proposed for its 26 6.7. Conclusion formation; this lens could have resulted either from a lee eddy eroding since the late winter, or from the fragmentation of the PGW outflow, at Ra's Al Hamra or south of Ra's Al Hadd, 15 or 10 days before the recording. This submesoscale lens was then observed between two strong mesoscale eddies during the Phys-Indien experiment, and having an elliptical structure. These mesoscale eddies deformed the lens, making it unsteady, and most likely, it rapidly disappeared afterwards.This mesoscale eddy shear/strain supports the formation of lenses from the coastal PGW outflow during the spring inter monsoon; this is supported by the repeated observations of such submesoscale patches of PGW, both with the SeaSoar and with the ARGO floats. But simultaneously the deformation field, induced by the mesoscale gyres, makes these lenses subject to break up, particularly out of the Sea Of Oman.

		Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring
	Month	January February March	April	May	June	July	August	September October November	December
	Density (kg.m 3 )	26.62	26.61	26.82	26.95	26.92	26.32	25.91	26.01	26.25	26.01	25.74	26.34
	Salinity (psu)	38.46	38.26	38.29	38.48	38.50	37.89	37.53	37.63	37.83	37.66	37.32	38.29
	Temperature ( o C)	22.79	22.31	21.66	21.70	21.87	22.33	22.80	22.74	22.42	22.78	22.84	23.32
	Thus, spring presents the most favourable conditions for PGW lens detection, with their ejection
	from the coastal outflow, and their advection around mesoscale eddies. During the summer
	monsoon, the PGW outflow is then expelled by the Ras Al Hadd jet, an intense mesoscale
	surface dipole, which may disrupt the PGW fragments shortly after their formation. In March
	2014, a second Phys-Indien took place around the Arabian Peninsula with a different mesoscale
	circulation, the study and comparision with spring 2011 situation of the PGW outflow forms will
	be the subject of a further study.
							27

TABLE 1 :

 1 Maximal PGW density, and associated salinity and temperature, before cascading in the Sea of Oman for each month extracted from the GDEM climatology.

	32

  . . . . . 47 Fig. 14. Sections across the lens off Ra's Al Hadd from surface down to 350 m depth. Measurements are : VM-ADCP and geostrophic velocities (from SeaSoar density), positive towards the west; Okubo-Weiss paramater calculated from the ARGO-GDEM density and velocity fields and Ertel Potential Vorticity from VM-ADCP and SeaSoar fields; s 0 potential density and spiciness. . . . . . . . . . . . . . . . . . . . . . . . 48 Topographic map of the Arabian Sea and Sea of Oman with the locations of interest. 35 Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring Location of the PhysIndien 2011 measurements. The red line follows the VM-ADCP and SeaSoar lines (with the exception of the Red Sea and Gulf of Aden where only the VM-ADCP was activated). The blue crosses and green circles represent the positions of XBT-XCTD casts and of CTDL-ADCP stations respectively. The black crosses represent the launch positions of floats (Surdrift and PROVOR).37 Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring Maps of the MADT (altimetric) anomaly averaged over each month, from November 2010 to June 2011.

		6.8. References 6.8. References 6.8. References Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring
			MADT anomaly in November 2010 MADT anomaly in March 2011 0.25			MADT anomaly in December 2010 0.2	0.25 Surface relative vorticity in March 2011 MADT anomaly in January 2011	0.25
		o N	o N					0.2		27 o N				27 o N					0.2		27 o N			0.2
								0.15				0.15						0.15					10	0.15
		o N						0.1		24 o N									0.1		24 o N			0.1
		o N o N o N o N o N o N					-0.2 -0.15 -0.1 -0.05 0 0.05	MADT anomaly [m]	15 o N 18 o N 21 o N		0.05 MADT anomaly [m] 0.05 0 0.1	21 o N 24 o N					-0.2 -0.15 -0.1 -0.05 0 0.05	MADT anomaly [m]	15 o N 18 o N 21 o N			0 5	-0.2 -0.15 -0.1 -0.05 0 [10 -6 s -1 ] 0.05	MADT anomaly [m]
	Map of salinity from the Argo floats at 250 m depth during a climatological month of March 0.2 Annual averaged and 2011 wind stress (N.m -2 ) Map of temperature from the Argo floats at 250 m depth during a climatological month of March 6 MADT anomaly in February 2011 MADT anomaly in March 2011 x 10 -7 Annual averaged and 2011 wind stress curl (N.m -3 ) Map of the Argo floats distribution at 250 m depth during a climatological month of March MADT anomaly in April 2011 54 o E 56 o E 58 o E 60 o E 62 o E 64 o E -0.25 54 o E 56 o E 58 o E 60 o E 62 o E 64 o E -0.25 54 o E 56 o E 58 o E 60 o E 62 o E 64 o E o N 0.1 18 N o -5	-0.25
	27 o N 0.18	o N o N			37		27 o N 0.25 0.2	Mean wind stress 2011 wind stress 27 o N		5 0.15			19.5 20		0.25 0.2	27 o N	27 o N			Mean wind stress curl 2011 wind stress curl -10	0.25 0.2
	0.16		o N					0.15					4	15 o N		19			0.15					0.15
	24 o N 0.14	o N	54 o E		56 o E	58 o E		24 o N 0.1 E o 60	62 o E	24 o N	64 o E	3 0.2	54 o E	18.5	0.1 56 o E	24 o N 58 o E 24 o N	60 o E	62 o E	64 o E	0.1
	21 o N 0.08 0.1 0.12	o N o N o N	Salinity [psu] Surface Okubo-Weiss field in March 2011 36.5 21 o N MADT anomaly [m] -0.05 0 0.05 21 o N		1 2			17 17.5 18	Temperature [ o C]	-0.05 0 0.05	21 o N MADT anomaly [m]	21 o N			Number of points -0.05 0.05 0	MADT anomaly [m]
	0.06	o N				36		-0.1		18 o N			0			16.5		-0.1		18 o N			-0.1
	18 o N 0.04	o N					18 o N -0.15					-1			16			-0.15	18 o N				-0.15
	0.02	o N	o N					-0.2		15 o N			-2			15.5		-0.2		15 o N			-0.2
	15 o N	54 o E	56 o E	58 o E	60 o E	62 o E	15 o N 64 o E -0.25		54 o E	56 o E	58 o E	60 o E	62 o E	15	64 o E	-0.25	15 o N	54 o E	56 o E	58 o E	60 o E	62 o E	64 o E	-0.25
	54 o E o N o N o N 54 o E 56 o E o N o N FIG. 1: 54 o E 58 o E 60 o E 56 o E 58 o E 62 o E 56 o E 58 o E 60 o E 62 o E 64 o E 35.5 FIG. 2: Maps of the salinity (left) and temperature (center) from GDEM modified by ARGO floats at 250 m depth for 64 o E MADT anomaly [m] -0.25 -0.2 -0.15 -0.1 -0.05 0 0.05 54 o E 56 o E 58 o E 60 o E 62 o E 64 o E 15 o N 18 o N 21 o N MADT anomaly [m] -0.25 -0.2 -0.15 -0.1 -0.05 0 0.05 60 o E 62 o E 64 o E 54 o E 56 o E 58 o E 60 o E 62 o E 64 o E 54 o E 56 o E 58 o E 60 o E 62 o E a climatological month of March; associated number of floats is indicated on the right-hand panel. 36 o E 42 o E 48 o E 54 o E 60 o E o N Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec -3 o N MADT anomaly in May 2011 0.1 0.15 24 o N 0.1 0.15 FIG. 3: Jan 0 0.25 MADT anomaly in June 2011 FIG. 4: Daily wind stress (left) and wind stress curl (right) in blue and their 15 day means in red, over a 2 o square region off Ra's Al Hadd. o N 0.2 0.2 27 o N 0.25 o N	64 o E
			34 36 38 FIG. 5: 39

  Eulerian volumic transport across the SeaSoar measurements in the Sea of Oman (top) and in the Arabian Sea (bottom). Arrows indicates the direction, the values are in Sverdrup and in color is indicated the maximum of salt in the PGW layer.FIG. 13: ARGO floats 2901370 (top) and 2901387 (bottom), trajectories (left) and salinity section (right) in the eastern and western Sea of Oman. Float 2901370 (up) enters the Sea of Oman with anticyclonic loops, is stopped by the front in March 2011 and then looped cyclonically until late May 2011, between A1 and A3. Float 2901387 looped cyclonically from March to May 2011 before moving northwest. Patches of salty PGW are observed, with the strongest in July when the float is found near the position of cascading PGW. 47 Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring Sections across the lens off Ra's Al Hadd from surface down to 350 m depth. Measurements are : VM-ADCP and geostrophic velocities (from SeaSoar density), positive towards the west; Okubo-Weiss paramater calculated from the ARGO-GDEM density and velocity fields and Ertel Potential Vorticity from VM-ADCP and SeaSoar fields; s 0 potential density and spiciness.

		6.8. References Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring 6.8. References 6.8. References
					Map of the temperature at the depth of maximal salinity -100 Volumic transport in Sverdrup in the Sea of Oman 22 Map of the PGW maximum salinity measured from SeaSoar Velocity section from VM-ADCP measurement 0.6 38 -50 0.8	38
																				0.4
		27 o N o N	Section across the Lens	Depth (m)	-250 -200 -150								27 o N					-0.2 0 0.2	37.8 [m/s]	37.8
		26 o N in the eastern Sea of Oman 21.8 -350 Trajectory of an ARGO float from December 2010 to May 2011 21.9 22 o N -300 o N	22.1	21 22.2	22.3	22.4	22.5	22.6	22.7	22.8	-0.8 -0.6 -0.4	37.6	37.6
		24 o N o N o N					-50					May 2011	20 Geostrophic velocity 24 o N Latitude ( o N)					0.8	37.4	37.4
	26 o N	21 o N 54 o E 56 o E 58 o E 60 o E 62 o E 64 o E 24 o N o N o N o N	Depth (m)	-350 -300 -250 -200 -150 -100	21.8	Temperature [ o C] 0.4 0.67 18 19 Feb 2011 0.21 0.64 0.14 0.28 0.53 0.14 0.04 Latitude ( o N) 0.53 21 o N 0.16 Mar 2011 -150 21.9 22 22.1 22.2 22.3 22.4 -100 Apr 2011 -50 Apr 2011 Salinity section of an ARGO float from December 2010 to May 2011 in the eastern Sea of Oman Salinity [psu] 36.6 36.8 37 37.2 0.48 36.8 22.5 22.6 22.7 22.8 -0.8 -0.6 Salinity [psu] -0.4 37 37.2 [m/s] -0.2 0 0.2 0.4 0.6	36.8 37 37.2 37.4
	24 o N	18 o N					-50 0					Feb 2011	18 o N 0.39 Okubo-Weiss field -200		0.94			0.3 0.4	36.6	36.4	36.6
	56 o E 22 o N	54 o E 15 o N	58 o E	56 o E	58 o E -350 60 o E 21.8 -300 -250 -200 -150 Depth (m) -100	60 o E 21.9	62 o E 22	62 o E 22.1	64 o E 22.2 16 Dec 2010 Jan 2011 17	54 o E Jan 2011 22.5 15 o N Dec 2010 22.3 22.4 -350 24 26 -300 1.03 -250	56 o E Feb 2011 22.6 22.7 58 o E Feb 2011 22.8 1.24	60 o E -0.4 Mar 2011 36.2 62 o E -0.3 -0.2 36.4 [10 -10 -0.1 0 s -2 ] 0.1 0.2	Apr 2011	64 o E	Apr 2011 Anticyclone A1 May 2011 Density σ 0 Anticyclone A3 Low spice front High spice front Between A1 and A3 36 36.2	36 36.2 36.4
	o N	18 20 22 Latitude ( o N) Ertel Potential Vorticity field 26.6 26.7 26.8 27 o N 58 o E 15 Aug 2011 -50 07 Jul 2011 26 Jul 2011 Volumic transport in Sverdrup in the Arabian Sea Map of the density σ 0 at the depth of maximal salinity -300 Trajectory of an ARGO float from March to August 2011 o N Map of the spice at -0.6 the depth of maximal salinity 56 o E 22 o N 36 in the western Sea of Oman 0 60 o E 24 o N 0.58 37 -0.4 -250 Salinity section of an ARGO float from March to August 2011 in the western Sea of Oman 27.5 28 -200 -150 [10 -12 -0.2 0 Depth (m) s -1 ] 0.2 -100 0.4 -50 0.6	38 38.2
	o N	o N o N			20 o N 22 o N Depth (m)	-350 -200 -150 -100 -50	21.8	21.9	22	22.1	35.8 21 o N 24 o N 22.4 1.52 N) o 14 16 22.3 σ 0 [kg.m -3 ] 26.3 26.4 26.5 16 Jun 2011 -100 22.2 Latitude ( 07 May 2011 -250 -200 27 May 2011 -150 σ 0 density section	22.5 1.02 6.43	36 22.6 1.93	36.2 22.7	22.8	36.4 25.5 26 26.5 27	[kg.m -3 ]	36.6 36.8	36.8	26.5 27	[kg.m -3 ]	37	37.2 37.8 37.6 36.8 37 37.2 37.4
	56 o E o N	FIG. 10: Potential temperature over salinity profiles (right) and their locations of interest (left); in the western Sea of Oman (up) and eastern Sea of Oman (down). Top panel : -Blue diamond : Profile out of the salty outflow; -Green cross : 58 o E lens; -Yellow circle : southern profile in the PGW outflow along the coastal slope; -100 27 -Grey/Black cross : 57.5 o E lens; o N 26.1 26 17 Apr 2011 -250 25 36.6 -300 -300 24.5 -350 26.2 18 o N 25.5 18 2.77 58 o E 60 o E 28 Mar 2011 28 Mar 2011 17 Apr 2011 07 May 2011 27 May 2011 16 Jun 2011 07 Jul 2011 26 Jul 2011 -350 21.8 21.9 22 22.1 22.2 22.3 22.4 22.5 22.6 22.7 24 22.8 Latitude ( o N) 62 o E Spice section o N 0.6 36.4 -50 3.23 28	15 Aug 2011	36.6 36.4 36.2
		54 o E o N		-Red circle : northern profile in the PGW outflow along the coastal slope; -Brown circle : cascading PGW outflow. Bottom panel : -Yellow circle : profile inside the anticyclone A3; 56 o E 58 o E 60 o E 62 o E 64 o E 25.9 26 54 o E 56 o E 58 o E 15 o N 16 o N 0.09 0.8 -300 2.38 0.38 -250 0.3 -200 -150 Depth (m)		60 o E 25 26	] [kg.m -3	62 o E 36.2	64 o E	25
		-350 -Blue diamond : Low spice profile; -Green diamond : High spice profile; FIG. 11: Scatter maps from the SeaSoar measurements displaying the maximal thermohaline characteristics of the Latitude ( o N) 0.05 21.8 21.9 22 22.1 22.2 22.3 22.4 22.5 22.6 22.7 22.8 24
		-Cyan cross : at the periphery of the anticyclones A1 and A3; -Red circle : profile inside the anticyclone A1. 44 Persian Gulf Water (for s 0 between 26 and 26.7). The variables are : temperature (top, left); salinity (top, right); s 0 at the maximal salinity depth (bottom, left); and spiciness at the same depth (bottom, right). 45 54 o E 56 o E 58 o E 60 o E 62 o E 64 o E 14 o N 36 FIG. 12: 46 FIG. 14: 48

  Chapter7. Evidence of Mediterranean Water dipole collision in the Gulf of Cadiz

	Journal of Geophysical Research: Oceans	10.1002/2014JC009972

L'H EGARET ET AL.

V C 2014. American Geophysical Union. All Rights Reserved.

3 respectively (Table

  Analysis of the experiments data showing the eddy interaction the recorded temperature did not vary substantially; we will use it here as a marker of this meddy (see Figure2a, the drift of the meddy center, calculated from the RF1 position records as in[START_REF] Paillet | Real-time tracking of a northern meddy in 1997-98[END_REF]).In February 2001, hydrological sections from the MedTop experiment revealed another dipole southwest of Portimão Canyon. The meddy was centered near 36 N, 8 45 0 W and the cyclone near 36 N, 8 15 0 W (seeFigure

	Journal of Geophysical Research: Oceans	10.1002/2014JC009972

L'H EGARET ET AL.

V C 2014. American Geophysical Union. All Rights Reserved. 4 7.5.

Table 1 .

 1 Table of Eddies Names, Period of First Measurement, Drifts, Interactions, and Floats Chapter7. Evidence of Mediterranean Water dipole collision in the Gulf of Cadiz

	Journal of Geophysical Research: Oceans	10.1002/2014JC009972
	Eddy	First Measure	Drift		Interaction	Floats
	M1	November 2000	East	Merged with M2	RF1
	C1	November 2000 (via altimetry)	East	Interacted with C2	None
	M2	February 2001	Southwest	Merged with M1	RF2 and RF3
						from April
	C2	February 2001	Southwest	Interacted with C1	None
	M3	June 2001	Rotates around C3	Results from the	SF
					merging of M1 and M2
	C3	June 2001	Stationary from April to June			None

L'H EGARET ET AL.

V C 2014. American Geophysical Union. All Rights Reserved.

5

  Chapter7. Evidence of Mediterranean Water dipole collision in the Gulf of Cadiz

	Journal of Geophysical Research: Oceans	10.1002/2014JC009972

L'H EGARET ET AL.

V C 2014. American Geophysical Union. All Rights Reserved. 9

Table 3 .

 3 Table of the Numerical Experiment Parameters for Collinear, Opposite, and Symmetric Dipoles

	Eddy	V 0	R	x 0	y 0
	A1	20.7	0.53	21.5	20.4
	C1	0.7	0.53	21.5	0.4
	A2	20.7	0.53	1.5	0.4 offset
	C2	0.7	0.53	1.5	20.4 offset

Table 4 .

 4 Table of the Numerical Experiment Parameters for Opposite, Symmetric Dipoles With a Finite Angle Between Orientations

	Eddy	V 0	R	x 0	y 0
	A1	21.0	0.53	22.0	0.8
	C1	1.0	0.53	21.4	1.4
	A2	21.0	0.53	1.4	1.4-offset
	C2	1.0	0.53	2.0	0.8-offset

L'H EGARET ET AL.

V C 2014. American Geophysical Union. All Rights Reserved.

Table 5 .

 5 Table of the Numerical Experiment Parameters for Opposite, Asymmetric Dipoles

	Journal of Geophysical Research: Oceans	10.1002/2014JC009972
	Eddy	V 0	R	x 0	y 0
	A1	20.6	0.53	22.0	0.8
	C1	1.0	0.53	21.4	1.4
	A2	21.0	0.53	1.4	1.4 offset
	C2	0.6	0.53	2.0	0.8 offset

L'H EGARET ET AL.

V C 2014. American Geophysical Union. All Rights Reserved.

named from the Greek mythology, where the Argonauts carried the quest of retrieving the Golden Fleece aboard the Argo, a ship with the gift of clairvoyance.

6.5. Vertical structure of mesoscale eddies in the region in Sping 2011, and relation to PGW

Chapter6. A submesoscale lens of Persian Gulf Water o↵ the Omani coast in Spring

Remerciements

Acknowledments collègues. Merci aussi aux Sarah, Thibault, Renaud pour leurs amitiés.

Enfin merci à Mélodie pour son amour et son soutien durant la dernière ligne droite. Et puis merci à Ilwë.

Acknowledgements. The authors are grateful to one referee whose fine remarks and suggestions considerably improved the presentation and contents of this paper. The authors thank the ARGO program (ARGO Regional Center for the Indian Ocean, INCOIS, Indian Government) for providing the hydrological data of the profiling floats, to the CNES/AVISO center (France) for the altimetric maps, and to the NGDC/MGG (USA) for the ETOPO2 bathymetry. Thanks are due to Amy Bower for fruitful discussions and suggestions. This work was supported by DGA under the "COMINO" and "SYNBIOS" projects.

Edited by: J. A. Johnson The publication of this article is financed by CNRS-INSU. www.ocean-sci.net/8/227/2012/ Ocean Sci., 8, 227-248, 2012 3.9. Appendix A : Climatological maps Acknowledgments The authors thank Universite de Bretagne Occidentale and DGA for their support under REI program COMINO and ASTRID program SYNBIOS, during the course of this work. They also thank Dr A. Bentamy (LOS, IFREMER) for providing wind data and CNES/AVISO for the altimetric data. The authors are grateful to the referees for their in-depth analysis of this work and for an essential suggestion which improved the quality of this paper.

Acknowledgements. The authors are grateful to Gilles Reverdin, Fabien Durand, the two referees and Matthew Hecht whose suggestions and remarks greatly improved the contents of this paper. This work was supported by DGA under the SYNBIOS project.

Edited by: M. Hecht Acknowledgments Thanks are due to SHOM (French Navy Hydrographic and Oceanographic Service) and to the Captain and Crew of R/V D'Entrecasteaux for data collection and processing. We acknowledge use of MedTop for data collection and processing. This study is part of the Semane research program funded by SHOM, DGA, IFREMER, and UBO. Portuguese Fundac¸ão para a Ciência e a Tecnologia is acknowledged for financial support of the MedTop Project.

Journal of Geophysical Research: Oceans

Appendix 1: Validation of methods for horizontal maps of hydrology and altimetry

Firstly, we present two temperature maps, for spring, over years 2002-2012; one obtained with our interpolation method and the other created directly with ODV ( [START_REF] Schlitzer | Ocean data view user's guide[END_REF]; see Fig. 16). The agreement is good, except west of 57.5 • E (see below for the error on hydrology).

Secondly, we show the map of significative points for hydrology obtained by interpolation/extrapolation of float data: this corresponds to the weight attributed to each point of the grid, related to the density of measurements in the region (the marked points on Fig. 16 for example). Below 5 points, in particular west of 57.5 • , where no float ever profiled, the interpolation is too poor to be trusted. Elsewhere, the number of points is sufficient, more specifically near Ras al Hamra where PGW ejection occurs.

Thirdly, we also show the error map for sea surface elevation (see Fig. 17): this error is maximum near the coasts, but at the location of the eddies that we describe in the text, the error is only ±6 cm. This must be compared with a difference of 20-25 cm in altimetry, between the crest and trough of the anticyclonic and cyclonic eddies considered. Therefore, the altimetric maps that we show have a 20-30 % error for eddies near a coast, and a smaller error in the deep ocean.

Chapter 5

Mesoscale variability in the Arabian Sea from HYCOM model results and observations : impact on the Persian Gulf Water path Contents Appendix A: Interaction of a long Rossby wave with a circular seamount

In this section, we provide a simple analytical calculation for the interaction of a long, zonal Rossby wave with a localized, circular (parabolic-shaped) seamount. We use the two-dimensional or barotropic vorticity equation on the beta plane with bottom topography:

Our basic flow is a zonal Rossby wave with wave number k

which is a solution of the nonlinear vorticity equation over a flat bottom with c = k 2 . We also recall that

Furthermore, we assume that the seamount height is small, r), and that it has a radius r 0 ⌧ 2⇡/k. Since the topographic term is small in the vorticity equation, we can look for a response of the flow in terms of a perturbation expansion (x, y, t) = 0 (x, t) + ✏ 1 (x, y, t).

(A4)

Thus we have, at order ✏,

To simplify this equation further, we assume that

), so that the last term is simply

or, more simply, a term in y exp(ik[x ct]) that is going to force the flow response in 1 . Thus we assume that

Then,

Therefore, the vorticity anomaly over the seamount is

and

which cancels out the second term of the time derivative. Therefore

or finally

The velocity associated with the zonal Rossby wave is meridional v 0 = @ x 0 = ikA 0 exp(ik[x ct]). Therefore

From this, one can see immediately that negative vorticity will be created by a northward motion of the particles (v 0 > 0) on the southern side of the seamount (y < 0, where the fluid columns will experience squeezing) and conversely on the northern side of the seamount. If the wavelength of the Rossby wave is long compared with the radius of the seamount, enough perturbation vorticity will build up on these two sides of the mount and a dipole will be formed that can then detach. An interesting feature of the water masses is observed around the strong mesoscale eddies, and lenses : injections of colder and fresher IOCW occur around them both in the Arabian Sea and in the Sea of Oman (see figure 8, cyan diamond profile). These structures present no front in density but a marked one in spiciness (see figure 7). Smith and Ferrari (2009) suggested that these filaments could here result from the stirring of PGW by the mesoscale eddies. The slope of these structures is compared with the f/N ratio, and with the strain over shear ratio induced by the eddies. The isospice slope varies from 0.3 to 0.4 %, f/N from 6 10 3 to 7 10 3 , and dU/dx/dU/dz is very close to the isospice slope. This suggests that stirring can form these lateral injections.

Furthermore, the presence of these colder and fresher injections below warmer and saltier PGW water favors double diffusion. This might be an explanation for the salty patch observed below the ASHSW layer on the southern profile (yellow diamonds in figure 8).

a. Submesoscale lens off Ra's Al Hadd

The lens section displayed on figure 6 corresponds to the measurements achieved between the 19 th and the 20 th March 2011. This section off Ra's Al Hadd extends between the southern cyclone C1 described earlier and the eastern part of anticyclone A3 at the mouth of the Sea Of Oman. The surface vorticity field indicates that the cyclone was more energetic than the anticyclone. This is also observed on the Okubo-Weiss field : the section is located in a region where deformation dominated rotation; indeed deformation reached there its maximal value observed in the Sea of Oman and Arabian Sea at this period. Note that the velocity induced by the mesoscale eddies was above 0.5 m.s 1 in the southern part of the section. 22 6.6. PGW characteristics, pathway and submesoscale structures A vertical profile (green cross profile on figure 8) and sections (see figure 14), indicate that this submesoscale lens has a density s 0 = 26.6 kg.m 3 , and a salinity above 37.3 psu; this salinity was observed only in the Sea of Oman at this period; therefore it is likely that this sea was the formation site of this lens. In L' Hégaret et al. (2015), several mechanisms leading to the formation of eddies containing PGW were cited. Hereafter, two most probable mechanisms are assessed.

In winter, the formation of coastal lee eddies is observed, downstream of Ra's Al Hamra; these eddies retain high salinity in their core before eroding during three months. This mechanism was observed in the high resolution, HYCOM simulation, and lee eddies were the only structure

Presentation of the article

The following article, published in July 2014 in the Journal of Geophysical Research, focus on the merging of two dipoles, containing both a Mediterranean Water, in the Gulf of Cadiz. For this, datasets from three di↵erent experiments were used, as satellite data and numerical simulations. This study, done during this thesis, used similar processes on the in-situ measurements and similar comparisons between floats seeded inside the mesoscale structures and the surface signature.

Beyond the data and methods used on this study, the analysis of the Mediterranean and Persian Gulf salty water mass pathways and characteristics allows to outline the differences between the two outflows and their propensities to destabilize and form lenses. where applicable (i.e., there is no coupling of the upper layer with the atmosphere, nor of the lower layer with the bottom). The coupling coefficients are F j;j21 5f 2 0 =ðg 0 j21=2 H j Þ; g 0 j21=2 5gðq j 2q j21 Þ=q 0 where f 0 5 8.34 3 10 25 s 21 is the Coriolis parameter at 35 N, H j is the thickness of layer j (H 1 5 600 m, H 2 5 900 m, H 3 5 2500 m). The beta value is b 5 1.86 3 10 211 m 21 s 21 . All these parameters are made dimensionless in the model.

RESEARCH ARTICLE

The stratification of this model is fitted to the region: the reduced gravities g 0 j21=2 are calculated such that the external deformation radius is infinite and the first two internal deformation radii are 30 and 15 km, respectively [see [START_REF] Carton | Vortex stability in a multi-layer quasi-geostrophic model: Application to Mediterranean Water eddies[END_REF]. We have Dq 1 q 0 5 10 23 and Dq 2 q 0 53310 24 .

The model performs spatial derivatives via a Galerkin projection on Fourier modes with truncation and a mixed Euler-leapfrog time stepping scheme is used.

The vorticity dissipation in the model is achieved with a harmonic and a biharmonic operator; the two viscosities are maintained to a minimum compatible with the absence of noise at small scales. The dimensional values are m 2 5 0.4 m 2 s 21 and m 4 5 3.2 10 7 m 4 s 21 , which are very low.

For simplicity, we confined the eddy velocity in the middle layer (note that potential vorticity thus exists in all layers):

where x 0 and y 0 are the coordinates of the eddy center. The Gaussian velocity profile is commonly used for MW eddies [see [START_REF] Carton | Vortex stability in a multi-layer quasi-geostrophic model: Application to Mediterranean Water eddies[END_REF]. Therefore, each eddy is characterized by four parameters V 0 , R, x 0 , y 0 . Note that the velocity maximum is

The Numerical Experiment Simulating the Observations

To reproduce the details of the interaction between the two dipoles, we started from the February 2001 situation. The previous stage of the evolution of M1-C1 has indeed been modeled before. This date corresponds to the formation of M2-C2 south of Portimão Canyon and thus to detailed measurements of this dipole. We do not have hydrological measurements of M1-C1 at the same date but we can estimate that its structure has not changed much since November 2000, but for a slight attenuation.

Also we center this situation in the middle of our domain, so that the origin is now at 36 N, 9 W. We locate M1-C1 using the Rafos floats trapped in M1, and we use the trajectory of M1 center (southward at that time) to infer that C1 lay east of M1 at that time (see Figures 2a and2b). The velocity of C1 is that used in [START_REF] Carton | Meddy dynamics and interaction with neighboring eddies southwest of Portugal: Observations and modeling[END_REF]. The velocity of M1 was chosen equal to 80% of its value in November 2000 to allow for a slight decay since its first observation.

Therefore, the parameters for this situation are, in dimensionless form:

Note that, for instance, V 0 5 0.93 corresponds to V max 5 0.4 m s 21 (and V 0 5 0.7-0.3 m s 21 ); R 5 0.707 leads to R max 5 20 km (Table 2).

In Figure 13, horizontal maps of potential vorticity in the middle layer are shown, zoomed in on the central region where interaction occurs (that is, in a 251 3 251 km). The initial position of M2-C2 is near Portimão Canyon, shifted here to the northeastern part of the subdomain. The initial position of M1-C1 is about 100 km further west and about 15 km north of the latter. Initially, C1 is east of M1 and C2 is southeast of M2. This corresponds to the February situation.

After 2 weeks, dipole M1-C1 has rotated cyclonically (C1 having a larger circulation than M1) and dipole M2-C2 has propagated southwestward. After 3 weeks, M2 is advected by C1, C2, and M2 and deformed by the flow field of these three eddies. M2 is considerably elongated. After 4 weeks, the merger of M1 and M2 takes place. A substantial mass transfer occurs from M2 to M1; but as it

Abstract

The Northwestern Indian Ocean is a region with a complex oceanic circulation. It is composed of basins with different topography. The Red Sea is crossed by an oceanic ridge whereas the Persian Gulf is a plateau with a maximum depth of 90 meters. These basins open on the Arabian Sea via the Gulf of Aden and Sea of Oman respectively. Furthermore, the region is dominated by the monsoon, strong winds over the ocean in winter and summer. These seasons are separated by inter monsoons, seasons where the winds lose intensity and main direction. Winds and high heat fluxes induce strong evaporation, mainly in the Red Sea and Persian Gulf. These highly saline water masses flow in the Arabian Sea with a high seasonal variability.

The Indian Ocean is thus a region of strong scientific and geopolitical interest, numerous oceanographic experiment and numerical simulations focus on the geophysical processes in the region. Although, the differences between each basin, the interactions between water masses and the strong seasonal variability lead these studies to focus on a specific period or phenomenon.

Recently, numerous studies emphasize on the Red Sea and the Gulf of Aden, where mesoscale eddies impact the circulation of the saline outflow. These structures form from westward propagating waves from the Indian coast, interacting and destabilizing through the interactions with the winds or eddies.

The aim of this thesis is to study the mesoscale circulation in the northen Arabian Sea and in the Sea of Oman, which phenomenon dominates the circulation, how do they form and what are their structures, how do they interact between them and with the monsoon variability ? In the same time, a study emphasizes the impacts of these structures in the outflows from the Persian Gulf and the formation of smaller scale structures.

To answer these questions, numerous data are used : in-situ measurements through oceanographic experiments and floats, satellite data and re-analysis from model, climatologies, but also results from realistic simulation (HYCOM) or idealized model to study processes. A first chapter focuses on the evolution and measurements of ARGO floats across the Gulf of Aden and the Arabian Sea. It points out strong correlations between the surface and deep circulation. Furthermore, highly saline waters are observed inside and at the periphery of mesoscale eddies with different characteristics and are advected by those. The following chapter focuses on the dynamic of these structures in the Sea of Oman. It revealed the presence of a dipole during the spring inter monsoon. It set up is linked to the end of the winter monsoon. The anticylonical part of the dipole, East of the basin, is reintensified during the inter monsoon through winds and structures propagating from the Indian coast. This dipole creates a strong deformation field along the coast of Oman, thus tearing the outflowing Persian Gulf Water in the northern part of the basin, with a recirculation East of it and a spreading in the Arabian Sea West of it.

In a third chapter, the use of a model and observations allowed to describe the mesoscale circulation in the northern Arabian Sea and in the Sea of Oman. The different characteristics of the outflowing Persian Gulf Water where studied for each season to complete the previous chapter's analysis.

Finally, a chapter focuses on the formation a smaller scale structures following the ejection of the Persian Gulf Water. During the Phys-Indien experiment a submesoscale lens where registered in the Arabian Sea. This chapter presents its characteristics and an estimation of the mechanisms leading to its formation and evolution.