# Ocean memory effect on the dynamics of coastal heavy precipitation preceded by a mistral event in the northwestern Mediterranean

Cindy Lebeaupin Brossier, Philippe Drobinski, Karine Béranger, Sophie Bastin, Françoise Orain

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Cindy Lebeaupin Brossier, Philippe Drobinski, Karine Béranger, Sophie Bastin, Françoise Orain. Ocean memory effect on the dynamics of coastal heavy precipitation preceded by a mistral event in the northwestern Mediterranean. Quarterly Journal of the Royal Meteorological Society, 2013, 139 (675), pp.1583-1597. 10.1002/qj.2049. hal-00769660

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# Ocean memory effect on the dynamics of coastal heavy precipitation preceded by a mistral event in the northwestern Mediterranean

Cindy Lebeaupin Brossier,ª,ᵇ* Philippe Drobinski,ª Karine Béranger,ᵇ Sophie Bastinᶜ and Francoise Orainᵈ

In this study, we examine the possible role of 3D oceanic circulation in ‘connecting’ a strong mistral windstorm in the northwestern Mediterranean with a heavy precipitating event in southern France in November 1999. These two events are both governed by intense air–sea exchanges and are only separated by a week. From an atmospheric point of view, these two events may be considered as fully uncorrelated and independent.

However, using the MORCE regional coupled system with WRF atmospheric model and NEMO-MED12 ocean model, we show that the oceanic cyclonic circulation is a key feature, which can trap strong modifications of the ocean mixed layer over few days. In our specific situation, the mistral (5–9 November 1999) induces a rapid mixed-layer cooling. This cold anomaly is trapped in the cyclonic gyre in the Gulf of Lions and maintained over a week. This directly affects the following heavy precipitating event (12–13 November 1999) with less energy supplied where the cold anomaly is maintained and larger evaporation near the coast. This modulates the rainfall amounts over land, with an increase in the eastern part and a decrease in the western part compared to an uncoupled simulation. These simulations highlight the key role of the general ocean circulation, and especially of the dynamic of the cyclonic gyre in the Gulf of Lions on the ocean memory in the area. The ocean maintains the temperature anomaly over typical periods of a few days, which thus enables some indirect links between two weather systems in which air–sea interactions are significant.

**Key Words:** CORDEX; heavy precipitation; HyMeX; Mediterranean Sea; mesoscale processes; mistral; ocean–atmosphere coupling

Received 19 December 2011; Revised 16 August 2012; Accepted 3 September 2012; Published online in Wiley Online Library 31 December 2012

**Citation:** Lebeaupin Brossier C, Drobinski P, Béranger K, Bastin S, Orain F. 2013. Ocean memory effect on the dynamics of coastal heavy precipitation preceded by a mistral event in the northwestern Mediterranean. Q. J. R. Meteorol. Soc. **139**: 1583–1597. DOI:10.1002/qj.2049

<footnote>ªLaboratoire de Météorologie Dynamique (LMD), Institut Pierre Simon Laplace, (CNRS/Ecole Polytechnique/ENS/UPMC), Palaiseau, France</footnote>

<footnote>ᵇEcole Nationale Supérieure de Techniques Avancées (ENSTA), ParisTech, Unité de Mécanique, Palaiseau, France</footnote>

<footnote>ᶜLaboratoire Atmosphères Milieux Observations Spatiales (LATMOS), Institut Pierre Simon Laplace, (CNRS/UVSQ/UPMC), Guyancourt, France</footnote>

<footnote>ᵈCentre de Météorologie Spatiale (Météo-France), Lannion, France</footnote>

<footnote>*Correspondence to: C. Lebeaupin Brossier, GAME/CNRM (CNRS/Météo-France), 42 avenue G. Coriolis, 31057 Toulouse CEDEX, France. E-mail: cindy.lebeaupin-brossier@meteo.fr</footnote>
# 1. Introduction

The Mediterranean basin is characterized by a semi-enclosed sea and a pronounced surrounding orography. The Mediterranean region is frequently submitted to intense weather events, which are characterized by their small scale and a high variability, and which imply coupled processes between ocean, atmosphere and land surfaces. In particular, the northwestern Mediterranean (NWM) area (Figure 1) around the Gulf of Lions (GoL) is vulnerable to strong winds (mistral and tramontane) enhanced and channelled by the surrounding topography and to heavy precipitation events (HPEs) triggered by moist flow from the sea. These two kinds of intense event imply fine (kilometric)-scale processes in the atmosphere and ocean (flow interactions with topography or bathymetry, low-level winds, atmospheric convection, upwelling, dense water formation and mesoscale ocean gyres, etc.) and strong air-sea exchanges.

Mistral and tramontane are northerly and northwesterly winds blowing in southeast France (Zechetto and De Biaso, 2007), channelled into the Rhone and Aude valleys (Figure 1). They transport cold continental air over sea (Drobinski *et al.*, 2001, 2005; Guénard *et al.*, 2005, 2006; Salameh *et al.*, 2007) and induce both intense momentum and heat exchanges at the air-sea interface, and thus ocean mixed layer cooling and mixing (Lebeaupin Brossier and Drobinski, 2009; Small *et al.*, 2012). HPEs often occur during fall when the air-sea thermal contrast is maximal. They are mainly driven by low-level wind convergence which extracts heat and moisture from sea and transports them over land and the surrounding topography (Ducrocq *et al.*, 2008; Nuissier *et al.*, 2008).

In this article, we investigate a sequence of a strong mistral events (5–9 November 1999) (Guénard *et al.*, 2006; Lebeaupin Brossier and Drobinski, 2009) preconditioned by cyclogenesis over the Ligurian Sea (Figure 1), followed by an HPE (12–14 November 1999) in the southeast of France (Ducrocq *et al.*, 2003; Lebeaupin *et al.*, 2006). During the HPE, strong winds (>20 m s⁻¹, Figure 2(a)) with gusts of nearly 40 m s⁻¹, were observed in the coastal area and the rainfall amounts exceeded 100 mm in 48 h over a large area in southeast France (Figure 2(b)) and reached a maximum value of 624 mm in 24 h in the town of Lézignan-Corbières (L-C in Figure 1). The flash flood and storm surge along the GoL coast that rapidly followed the event led to more than 30 fatalities.

These two events were separated by about a week, which is already a long time-scale for mid latitudes when compared to the typical synoptic weather time-scales of few days. From an atmospheric point of view, these two events may be considered fully uncorrelated and independent.

The ocean is well known as a major component of the climate system. Indeed, owing to its large thermal capacity, the ocean keeps the integrated signature of atmospheric forcing and thus induces significant feedbacks in the integrated system at interannual to decadal time-scales. At shorter time-scales, i.e. synoptic weather time-scales (corresponding to a few days), the role of ocean ‘memory’ on the weather system predictability is rarely considered. Nevertheless, even at these short time-scales the processes in the ocean mixed layer and at its interfaces are crucial to fully represent air-sea interactions, especially during high-wind situations. In particular, ocean-atmosphere coupled models already show improvements in the representation of

tropical cyclones, with in particular the major role of cold-water entrainment across the thermocline, on sea surface cooling (Price, 1981; Vincent *et al.*, 2012) and consequently on tropical cyclone intensity (Bender *et al.*, 2007).

This work is an original investigation of the ocean memory effect considering 3D ocean processes and coupling at mesoscale during a sequence of two different kinds of intense weather events in the NWM region. In this article, we investigate the role that the Mediterranean Sea may play in linking the two meteorological intense events. We focus on the November 1999 sequence, with first a strong mistral event, which was then followed by an HPE. Previous numerical experiments on these specific cases, with 0D (slab ocean model) or 1D (independent water columns) ocean coupling, show that the ocean response is intense and immediate (Lebeaupin Brossier and Drobinski, 2009; Lebeaupin Brossier *et al.*, 2009). The signature of the intense exchanges during the mistral takes the form of ocean mixed layer cooling and deepening (Lebeaupin Brossier and Drobinski, 2009). For HPE, the ocean state, in particular the local heat content, plays a significant role on the initiation and feeding of the convective system (Lebeaupin *et al.*, 2006). The southeasterly low-level jet over the sea is also a key factor in the extraction of ocean energy and its transport. Owing to strong momentum and water fluxes, the thermohaline characteristics are rapidly modified, producing strong horizontal temperature and salinity gradients in the GoL (Lebeaupin Brossier *et al.*, 2009).

All the ingredients are thus present which justify such a study. However, in previous studies the short-range simulations (18–96 h) with 0D or 1D coupling only investigated *immediate and local* ocean-atmosphere feedbacks. The key issue addressed in this article is the possible role of 3D oceanic circulation on the persistence over a longer period of time of significant coupled processes which may indirectly link these two weather events.

The next section details the numerical design. Sections 3 and 4 describe the representation of the mistral event and the HPE, respectively, as well as their sensitivity to ocean-atmosphere coupling. Section 5 discusses the role of the ocean circulation and thermal evolution, and specifically its memory effect, on the link between the HPE and the mistral event. This last section also concludes the study and proposes a possible follow-up.

# 2. Experimental design

We developed a regional two-way air-sea coupled system over the Mediterranean basin within the MORCE (Model of the Regional Coupled Earth system) platform (Drobinski *et al.*, 2012). MORCE has been designed to investigate the role of coupled processes on the regional climate of particularly vulnerable areas. The MORCE system is used in the *Hydrological Cycle in the Mediterranean Experiment* (HyMex) (Drobinski *et al.*, 2009, 2010) and the *Coordinated Downscaling Experiment* (CORDEX) of the World Climate Research Program (WCRP) (Giorgi *et al.*, 2009).

## 2.1. *The atmospheric model*

The atmospheric model within the MORCE system is the *Weather Research and Forecasting* (WRF) model of the
<figure><img src="image_3.png" /><figcaption>**Figure 1.** The northwestern Mediterranean region (topography with contours every 500 m). L-C indicates Lézignan-Corbières town location (43.1°N, 2.45°E).</figcaption></figure>

National Center for Atmospheric Research (NCAR) (Ska-
marock *et al.*, 2008). The domain covers the Mediterranean
basin with a horizontal resolution of 20 km. It has 28 sigma
levels in the vertical. Initial and lateral conditions are taken
from the European Centre for Medium-Range Weather
Forecasts (ECMWF) ERA-interim reanalysis (Simons *et al.*,
2007) provided every 6 h with a 0.75° resolution. More-
over, to avoid unrealistic departure from the driving fields,
indiscriminate nudging is applied with a coefficient of
$5 \times 10^{-5} \text{ s}^{-1}$ for temperature, humidity and velocity com-
ponents above the planetary boundary layer (Stauffer and
Seaman, 1990; Salameh *et al.*, 2010). A complete set of
physics parametrizations is used with the WRF Single-
Moment 5-class microphysical scheme (Hong *et al.*, 2004),
the new Kain-Fritsch convection scheme (Kain, 2004), the
Yonsei University (YSU) planetary boundary layer (PBL)
scheme (Noh *et al.*, 2003) and a parametrization based on
the similarity theory (Monin and Obukhov, 1954) for tur-
bulent fluxes. The radiative scheme is based on the Rapid
Radiative Transfer Model (RRTM) (Mlawer *et al.*, 1997)
and the Dudhia (1989) parametrization for long-wave and
short-wave radiation, respectively. For the land surface, the
thermal diffusion scheme with five layers is used (Eckel,
2002).

## 2.2. *The ocean model*

The ocean model of MORCE is *Nucleus for European Modelling of the Ocean* (NEMO) (Madec, 2008). It is used in a regional eddy-resolving Mediterranean configuration MED12 (Lebeaupin Brossier *et al.*, 2011) with a 1/12° horizontal resolution (~7 km). In the vertical, MED12 has 50 stretched z-levels from (Δz = 1 m at the surface to Δz = 450 m at the bottom, with 35 levels in the first 1000 m). The initial conditions for 3D potential temperature and salinity fields are provided by the MODB4 climatology

(Brankart and Brasseur, 1998) except in the Atlantic zone
between 11°W and 5.5°W, where the Levitus *et al.* (2005)
climatology is applied. In this area, a 3D relaxation to this
monthly climatology is applied during simulations. The
runoffs and Black Sea water input are prescribed from a
climatology as surface freshwater gains, at the mouths of
the 33 main rivers or at the Dardanelles Strait. Elsewhere,
smaller river runoffs are gathered and put as a homogeneous
coastal runoff around the Mediterranean Sea (Beuvier *et al.*,
2010).

Horizontal eddy diffusivity is fixed at $K_h = 60 \text{ m}^2 \text{ s}^{-1}$ for tracers using a Laplacian operator (the diffusion is applied along iso-neutral surfaces) and the horizontal viscosity coefficient is fixed at $-1.25 \times 10^{10} \text{ m}^4 \text{ s}^{-2}$ for the dynamics (momentum) using a bi-Laplacian operator. The Total Variance Diminishing (TVD) scheme is used for tracer advection in order to conserve energy and enstrophy (Lévy *et al.*, 2001; Barnier *et al.*, 2006). Vertical diffusion is performed by the standard turbulent kinetic energy (TKE) model of NEMO (Blanke and Delecluse, 1993) and in case of instabilities a higher vertical diffusivity coefficient $K_z$ of $10 \text{ m}^2 \text{ s}^{-1}$ is used (Barnier *et al.*, 2006). A no-slip lateral boundary condition is applied and bottom friction is represented by a quadratic function with a coefficient depending on the 2D mean tidal energy (Lyard *et al.*, 2006; Beuvier *et al.*, 2012). The volume is finally conserved through a monthly damping of the sea surface height (SSH) between 11°W and 7.5°W towards an SSH dataset built from the GLORYS-1 experiment (Ferry *et al.*, 2010) and a former MED12 simulation (Beuvier *et al.*, 2012).

## 2.3. *Numerical experiments*

Control simulation (CTL) is the downscaling of the ERA-
interim reanalyses obtained with WRF alone. Sea-surface
<figure><img src="image_4.png" /><figcaption>**Figure 2.** Observations of the HPE: (a) daily 10 m winds from the Quickscat/SeaWinds scatterometer on 12 November 1999 and (b) 48 h accumulated rainfall (mm) between 12 and 14 November 0000 UTC in southeast France estimated in the SAFRAN gridded product.</figcaption></figure>

temperature (SST) is thus prescribed from ERA-interim and is updated daily.

The coupled simulation (named CPL experiment) runs with two-way interactive exchanges between the two compartment-models managed by the OASIS coupler (Valcke, 2006). The exchanged variables are the SST and the heat (solar Qs and non-solar Qns), water (Fw = E - P) and momentum (τ) fluxes. The coupling frequency is 3 h. The coupler uses a bilinear method to interpolate the ocean grid towards the atmospheric grid and vice versa.

Both simulations run from January 1989 to December 2008. CPL starts with an ocean at rest. In this study, we only focus on the NWM area at the beginning of November 1999, i.e. far from any drift.

### 3. The 5–9 November mistral event

A mistral event occurred between 5 and 9 November 1999 (Guénard *et al.*, 2006; Lebeaupin Brossier and Drobinski, 2009), induced by the passage of a cut-off low over the North Sea associated with a marked upper westerly flow blowing at 40 m s⁻¹ at 500 hPa at 0000 UTC on 6 November. A marked upper northwesterly flow behind the cold front is associated with the cyclone. After 0000 UTC on 7 November, the Genoa cyclogenesis is triggered in the lee of the Alps. The presence of the Azores ridge over western France maintains a moderate northwesterly wind at 500 hPa (around 20 m s⁻¹). At 0000 UTC on 8 November, the Genoa cyclone moves southeastward and the 500 hPa winds veer southward
<figure><img src="image_5.png" /><figcaption>**Figure 3.** Wind speed (ff in m s⁻¹, left panels) and direction (dd, right panels) in CTL and CPL experiments during the 5–14 November period over (a) the Gulf of Lions (5.0°E, 42.4°N), (b) Ligurian Sea (7.8°E, 43.1°N) and (c) Tyrrhenian Sea (13.0°E, 39.5°N) (the locations of the three grid points considered are indicated by a pink 'x' in Figure 1).</figcaption></figure>

over the northwestern Mediterranean until 0000 UTC on 9 November.

These upper-level conditions are accurately simulated by CTL and CPL experiments (not shown). Near the surface, the mistral onset occurs at around 0600 UTC on 6 November and veers progressively from the northwest to the north (Figure 3). The mistral peaks at 0600 UTC on 7 November with intensities over 20 m s⁻¹ and is quite steady for 24 h. Then, its strength slowly decreases on 9 November. Compared to CTL, CPL shows a slight decrease in wind speed, an increase in sea-level pressure and a small shift in the Genoa cyclone displacement towards the Tyrrhenian Sea, which slightly modifies the wind direction (Figure 3).

On 5 November, before the mistral event, the SST in CPL (Figure 4(e)) shows the presence of relatively cold water (about 17°C) in the GoL, whereas SST is about 19°C in CTL (Figure 4(c)). The CPL SST field is in good agreement with weekly satellite data (Figure 4(a)) from the French Center of Space Meteorology (CMS), obtained from the *Advanced Very High-Resolution Radiometer* (AVHRR) with 10 km resolution. This cold anomaly is the signature of the well-known cyclonic circulation in the GoL (Schott *et al.*, 1996)

induced by the winds and bathymetry (Madec *et al.*, 1996). Two other cold anomalies linked to cyclonic circulation in the Ligurian Sea (Ligurian gyre) and in the Tyrrhenian Sea (Bonifacio gyre) are well represented in CPL, and, on the contrary, are absent in CTL (ERA-interim) SST. The correlation values between the simulated and satellite SST fields on 5 November (Figure 4(a,c,e)) are 0.65 and 0.70 for CTL and CPL, respectively. The covariance between satellite and CPL SST fields in the NWM is 0.83 on 5 November, indicating good spatial correspondence between the two fields.

During the mistral period, strong wind and evaporation take place. A large cooling of the sea surface affects the NWM area. In CTL, the ERA-interim SST shows a global cooling in the area between 5 and 12 November 1999 (−0.2°C per day, but it is still very smoothed (Figure 4(d)) and thus filters out fine-scale cold anomalies evidenced by satellite dataset (Figure 4(b)). In CPL, from 5 to 9 November, the temperature decreases on average by 0.2–0.3°C per day but locally more. The SST decreased to about 14°C in the centre of the GoL (5°E–42°N, Figures 4(f) and 5), in very good agreement with satellite data (≈14°C) compared to CTL
<figure><img src="image_6.png" /><figcaption>**Figure 4.** SST fields (°C) over the NWE on 5 (left column) and 12 November 1999 (right column) for satellite product (AVHRR, top), CTL (middle) and CPL (bottom). The covariance and correlation with satellite data and the means of each dataset over the 0–12°N; 39.5–45°N box are indicated at the top of each panel. (g) SST differences (°C) between CPL and CTL (as seen by the atmospheric model) on 12 November 1999.</figcaption></figure>
<figure><img src="image_7.png" /><figcaption>**Figure 5.** CPL experiment: temperature (colour, °C) and salinity (contour, psu); section along the 5°E longitude (yellow dashed line in Figure 1: (a) at 1200 UTC on 5 November 1999 and (b) at 1200 UTC on 12 November 1999.</figcaption></figure>

(≈17.5°C in GoL). For both simulations, the correlation for SST is 0.73 on 12 November, but CPL shows the highest covariance score (0.96), which confirms the very good agreement in terms of spatial structures.

Figure 5 presents temperature and salinity section in the NWM (see Figure 1 for location) before the Mistral event (5 November) and one week later (12 November). It shows that, during the mistral event, salinity also increases due to strong evaporation (Figure 5). The ocean mixed layer cools and deepens (Figure 6(a)), in particular in the coastal area due to upwelling. Cold and dense water is also formed over the shelf during the mistral and tramontane event associated with a large Ekman pumping velocity (Figure 6(b)) and overflows. The event induces an increase by 100–125 m in the mixed-layer depth between 5 and 9 November along the Catalonian coast (3°E, 41–42.5°N). Inside the GoL gyre (4.8°E, 42.4°N), a strong cooling and deepening also take

place (Figures 4(g) and 5). In addition to the atmospheric contribution, the cold and salty anomaly in the GoL is enhanced by the entrainment of cold water (<14°C) initially present at 40 m depth (Figure 5(a)), i.e. at the ocean mixed layer base. Note finally that the cyclonic circulation together with the Northern Current bring warmer surface water near the French coast and over the GoL shelf (Figure 4(b,f,g)).

**4. The 12–13 November heavy precipitation event (HPE: ‘Aude case’)**

After the cyclone, which is at the origin of the mistral, left the NWM area, an upper-level low-pressure area settles down on 12–13 November 1999 over Spain and southern France, inducing a southerly flow from North Africa to southern France associated with high conditional convective instability. Inside the warm air mass advected northwards,
<figure><img src="image_8.png" /><figcaption>**Figure 6.** Mistral event in CPL: (a) ocean mixed layer depth differences (m) between 9 and 5 November 1999. (b) Ekman Pumping velocity $w_{Ek} = \bar{k} \cdot (\bar{\nabla} \wedge (\frac{\bar{k}}{\rho f}))$ (in $10^{-3} \text{ m s}^{-1}$, with $\bar{k}$ the vertical unit vector) on 7 November 1999 (i.e. during mistral maximum intensity).</figcaption></figure>

a surface low forms and accelerates the low-level easterly to southeasterly jet over the GoL. The sustained wind reaches $25 \text{ m s}^{-1}$ inside the moist low-level jet, and is accelerated by the deflection of the flow by the southern slope of the Alps (Figure 2(a)). Convection organizes in a quasi-stationary mesoscale convective system that stays over southern France from 12 to 14 November and generates intense precipitation (Figure 2(b)) organized in a narrow quasi-stationary line (Ducrocq *et al.*, 2002, 2003).

Figure 7 shows the low-level environment (evaporation and wind) in the two numerical experiments. The low-level jet is well reproduced compared to Quickscat satellite

data (Figures 2(a) and 7), and is associated with strong evaporation (LE between 200 and $500 \text{ W m}^{-2}$) especially near the French coast. The low-level jet thus supplies the atmospheric boundary layer with heat and moisture that it has extracted from the ocean. Indeed, the strong wind and evaporation again decrease the temperature in the area (and increase the salinity). High values of convective available potential energy (CAPE) are also found and the horizontal moisture convergence is maximal along the French coast (not shown). The moist and warm air tongue follows the southeastern branch of the low-level jet. Comparing CTL and CPL experiments, we find a decrease in wind speed
<figure><img src="image_9.png" /><figcaption>**Figure 7.** Latent heat flux (W m⁻²) and 10 m wind (m s⁻¹) at 0000 UTC on 12 November 1999 in (a) CTL and (b) CPL (light-blue lines indicate the 10 m wind speed (m s⁻¹) and dark-blue line indicates the cold anomaly (SST under 16°C) in CPL); (c) differences in latent heat flux (colour, W m⁻²) and 10 m wind speed (contours every 2 m s⁻¹, blue for decrease and red for increase in CPL).</figcaption></figure>
<figure><img src="image_10.png" /><figcaption>**Figure 8.** Difference in horizontal moisture convergence ($Q_{cvg} = -\vec{\nabla} \cdot (q\bar{v})$ in $10^{-6} \text{ kg kg}^{-1} \text{ s}^{-1}$) between CPL and CTL, 12 November 1999, 0000 UTC.</figcaption></figure>

above the cold anomaly in the GoL by 2–8 m s⁻¹ locally on 11 and 12 November (Figure 3). The latent heat flux patterns show a direct response to the SST fields. The evaporation is reduced over the cold anomalies from 200 W m⁻² in CTL to 100 W m⁻² in CPL at 5°E, 42°N, for instance (Figure 7). Conversely, it is stronger near the coasts (550 W m⁻² in CPL compared to 500 W m⁻² in CTL). The moisture convergence is significantly reduced in CPL close to the cold anomaly (Figures 4(g) and 8) due to less evaporation with modifications of the low-level wind (Figure 7(c)), whereas it increases in the coastal area, where SST (evaporation) is higher (Figure 7). The CAPE maximum at 6°E, 40.25°N is also larger in CPL (not shown).

In the two simulations, strong precipitation occurs from 12 to 14 November 1999 (Figure 9). The location and the 'Y' structure of the mesoscale convective system are accurately represented compared to the SAFRAN precipitation analyses (Habets et al., 2008) (Figure 2(b)) and to very high-resolution (2.5 km) numerical simulations (Ducrocq et al., 2002, 2003; Lebeaupin Brossier et al., 2009). An eastward shift of the heavy precipitating band (between Massif Central and Pyrenees in the observations; Figure 2(b)) of 40–80 km is found. Rainfall intensity is also underestimated in the two simulations. The maximum 48 h accumulated rainfall is 310 mm in CTL and 321 mm in CPL (compared to 624 mm observed). The maximum location is unchanged at 2.9°E, 43.6°N (Figure 9) and corresponds well to the observed maximum location (Figure 2(b)).

The coupling has a direct impact on precipitation over sea with less precipitation over the cold anomaly area simulated in CPL, and more precipitation where the SST is warmer (Figure 9). Over land, the coupling has an indirect effect and modulates the energy supply by the low-level jet. The cold anomaly (associated with a decrease in the wind speed) leads to a drastic decrease of the heat and moisture supply from the south in CPL. Conversely, the warmer SST near the coast induces an increase in evaporation and brings more heat and

moisture. Finally, the precipitation amount simulated over land in CPL is lower over the western part of the precipitating area with respect to the CTL simulation (Figure 9(b)). This is directly linked to the southeastern branch (offshore) of the low-level jet flowing over the cold anomaly. On the contrary, the precipitation amount simulated in CPL is larger over the eastern part associated with the northern branch (coastal) of the low-level jet.

Finally, the SSH field at 11 and 12 November in CPL (Figure 10) indicates the rapid sea-level rise of 25 cm over the shelf of the GoL close by the HPE location. This ocean response was observed and highly contributed to amplify the floods in the Aude region that followed the intense precipitation event.

## 5. Discussion on the ocean memory effect

From Section 4 it appears that the HPE is impacted by the presence of the cold anomaly in the GoL, which has its origin in the strong preceding mistral event. The cold anomaly persists several days after the mistral event and directly modifies the heat and moisture supply under the low-level jet for the following HPE: less energy is supplied from the south, where the cold anomaly is maintained, whereas larger evaporation occurs near the coast, where the SST is warmer in CPL. This does not change the chronology of the precipitating event, but it modulates the rainfall amounts, even over land, with an increase in the eastern part and a decrease in the western part. There is thus an indirect link between the mistral event and the HPE through the 'memory' held in the ocean.

Comparing in this study the 3D-coupled and uncoupled simulations during the mistral, we found small modifications of wind intensity but a small shift of the Genoa cyclone motion. This slightly changes the wind direction, in agreement with Lebeaupin Brossier and Drobinski (2009) comparing 0D-coupled/uncoupled experiments. Over the
<figure><img src="image_11.png" /><figcaption>**Figure 9.** 48 h accumulated rainfall amount (mm, between 12 and 14 November 1999) in (a) CTL (colour) and (b) CPL (green contours); colours in (b) indicate the differences between the two fields (mm).</figcaption></figure>

cold anomaly in CPL, the wind speed is lower during the HPE with respect to the uncoupled run, because of the stabilizing effect which reduces wind intensity and variability. This result is in agreement with Pullen *et al.* (2006) and Small *et al.* (2012), who use the 3D coupled model COAMPS. Finally, the SST field simulated with MORCE remains very accurate over the whole investigated period (Figure 4), which is a key added-value of the 3D coupled run.

In fact, one limit of the experiments presented here and designed for long term (20 years) is that, considering the 20 km resolution of the atmospheric model, the cold ocean anomaly in the GoL is quite small (100 × 100 km², i.e. 5 × 5 grid points in the atmospheric model). Moreover, the 3-hourly coupling frequency does not allow complete representation of the mistral effects on the ocean, considering the higher frequency of the wind with gusts, and thus the strong variations in wind intensity that could occur in less than 5 min. The SST is also daily updated in the uncoupled run and, as ERA-interim SST includes observations, it contains an integrated signature at coarse resolution of the previous air-sea interactions. Nevertheless, the use of a prescribed SST in CTL experiment (in particular with a coarse resolution like ERA-interim) does not allow

the investigation of the full ocean processes involved in the response to intense weather events and their feedbacks.

To further investigate the processes involved in this memory effect, we examine the various features obtained with the 3D coupling (CPL simulation), also comparatively to former 1D/0D air-sea coupled experiments. In all the coupled simulations, an intense and immediate sea surface cooling under the 5–9 November 1999 mistral event is captured on average. In the full 0D coupled simulation (for further details see Lebeaupin Brossier and Drobinski, 2009), the cooling is –0.15°C per day on average in the NWM, whereas it is –0.22°C per day in CPL. Even if these mean values are close, in the latter simulation (CPL) air-sea fluxes and the local circulation produce a remarkably accurate SST field in terms of intensity and spatial distribution, with respect to the observations. The same conclusions are obtained during the HPE: the SST also decreases immediately in the GoL by about –0.2°C in the 3D-coupled model versus –0.1°C in a 0D-coupled model (Lebeaupin *et al.*, 2006) and –0.3°C in a 1D-coupled model (Lebeaupin-Brossier *et al.*, 2009). However, the location is better reproduced with 3D ocean coupling.

Indeed, a 0D ocean model only simulates the trend of the thermal characteristics of the ocean mixed layer due to the atmospheric fluxes (Lebeaupin *et al.*, 2006). A 1D
<figure><img src="image_12.png" /><figcaption>**Figure 10.** Sea Surface Height (cm) and 30 m depth currents (m s⁻¹): (a) at 1200 UTC on 11 November 1999 and (b) at 1200 UTC on 12 November 1999 in CPL.</figcaption></figure>
<figure><img src="image_13.png" /><figcaption>**Figure 11.** CPL experiment: trends time series of the mixed-layer temperature and contributing terms (°C per day): (left panels) in the centre of GoL gyre (red box) and (right panels) west of Sardinia (green box). The boxes are indicated in Figure 1.</figcaption></figure>

ocean model explicitly represents, in addition to 0D ocean models, entrainment and vertical advection and diffusion (Lebeaupin Brossier et al., 2009). In these models, the horizontal current is generally diagnosed from the wind stress as the mean Ekman current in the mixed layer, but does not play any role on the ocean dynamics. The ocean thermal response in these models is thus only local on the vertical and the response time to the surface forcing could be estimated as the ratio of the squared length scale to a diffusion scale: $h^2/K_z$, where $h$ is the ocean mixed layer depth (~50 m in November) and $K_z$ the vertical eddy diffusivity coefficient (enhanced to $10\ m^2\ s^{-1}$ in case of instabilities). The typical time response is thus $h^2/K_z \simeq 4$ min. Consequently, any ocean anomaly is rapidly forgotten if the atmospheric forcing is unsteady. The local 'vertical' memory is induced by entrainment (if any) that propagates an anomaly towards intermediate and deep oceanic layers. With an entrainment speed of $w_e = \frac{dh}{dt} = 5 \times 10^{-5} \ m\ s^{-1}$ (~12 m in 3 days; see Lebeaupin Brossier and Drobinski, 2009), the memory time could be considered as $h/w_e \sim 12$ days.

The benefit of the 3D ocean model is clearly the general (horizontal) circulation, with advection and gyre-induced vertical structures.

Figure 11 presents the contributing terms to the mixed-layer thermal trend in the centre of the GoL cyclonic gyre and outside (west of Sardinia) in the 3D coupled experiment (see boxes in Figure 1). During the sequence, the cooling due to forcing and diffusion are persistent and almost the same everywhere in the NWM (Figure 11). Outside the cyclonic gyre, the cooling by entrainment is moderate during the mistral. Moreover, west of Sardinia, the advection of warmer water from the south after the mistral event almost compensates the cooling induced by the surface forcing and entrainment. In the GoL cyclonic gyre, the ocean mixed-layer cooling on 7 November in the cyclonic GoL gyre is mainly due to entrainment (−0.75°C) and atmospheric forcing (−0.3°C) (Figure 11(a)). Between

the two events, the cold anomaly in the GoL is maintained by advection and entrainment, first triggered by the strong mistral stress and mainly controlled by the gyre dynamics the days after. Indeed, in the NWM, the GoL cyclonic gyre is persistent (Figure 10(a)), enhanced by frequent mistral and bottom topographic trapping (Madec et al., 1996). The wind induces, in addition to ocean mixed-layer cooling and deepening, both a deepening of the SSH (Béranger et al., 2010) and a doming of the isopycnes (Marshall and Schott, 1999). The doming of the isotherms (shown in Figure 5) evidences the presence of cold water (<14°C) close to the surface. Moreover, the doming is already present on 5 November, before the mistral. Indeed, this barotropic structure progressively appears under strong winds since the beginning of September (not shown), as strong wind events, frequently encountered in fall, reinforce the cyclonic horizontal circulation (Béranger et al., 2010). Thus the complete set of 3D processes (meaning horizontal and vertical (1D) processes) is necessary to produce and to maintain the GoL cyclonic gyre. This very stable 3D ocean structure traps and isolates the water masses at an intermediate level. The trapping of this cold water by the GoL cyclonic gyre prevents the vanishing of the cold anomaly. It induces a weak stratification, while the isolation favours steady ocean heat loss and amplifies the cold anomaly intensity and persistence.

This finally counteracts the destabilizing effects of baroclinic structures brought by the Northern Current with the horizontal advection (with a typical time-scale of $L/U \sim 6$ h to 1 day, with $L \sim 10$ km the Rossby deformation radius in the Mediterranean and $U = 0.1–0.5 \ m\ s^{-1}$; see Lebeaupin-Brossier et al., 2009, 2011) or the smoothing effect due to lateral eddy diffusion (Madec et al., 1991) (+0.02°C during the period (Figure 11(a)); with a typical time-scale of $L^2/K_h \simeq 19$ days, with $L \sim 10$ km and $K_h \sim 60 \ m^2\ s^{-1}$ the horizontal eddy diffusivity). Since such trapping cannot be captured by a local grid-point coupling (0D or 1D), there is clearly added value for a
3D ocean–atmosphere coupling with realistic general ocean circulation and air–sea interactions, even for short-living weather systems.

In the CPL experiment, the ocean maintains the atmospheric signal in the form of strong modifications of the thermodynamic content of the ocean mixed layer. This memory of the ocean mixed layer could typically last over periods of several days (1–3 weeks) and finally enables synoptic weather systems to indirectly affect the subsequent storms. Consequently, taking into account the air–sea 3D coupled processes has a significant role and clearly affects the atmosphere dynamics in the case of successive weather systems. This result is of particular interest for high-impact weather systems such as windstorms or heavy precipitation.

Based on this result, a natural follow-up of this study is to fully analyse the 20-year uncoupled and coupled simulations, to evaluate the occurrence of such indirectly linked synoptic weather phenomena, and to further examine, using 3D-ocean heat budget analysis, the main governing factors of the ocean memory effect on the atmosphere at synoptic weather time-scales.

## Acknowledgements

This work is a contribution to the MORCE-MED project funded by the GIS ‘Climat, Environnement et Société’ and was supported by the IPSL group for regional climate and environmental studies. This work also contributes to the HyMeX program (HYdrological cycle in The Mediterranean EXperiment) through INSU-MISTRALS support and the Med-CORDEX program (COordinated Regional climate Downscaling EXperiment –Mediterranean region), with granted access to the HPC resources of IDRIS (under allocation i2011010227). The authors are very grateful to Marc Stéfanon (LMD), Jonathan Beuvier, Guillaume Samson (Mercator-Océan), Sébastien Masson, Gurvan Madec (LOCEAN), Sophie Valcke, Laure Coquart and Eric Maisonnave (CERFACS) for their useful collaboration in the development of the regional WRF-OASIS-NEMO system. The authors also thank Hervé Roquet (CMS/Météo-France), who provided SST satellite data, and the Climserv team from IPSL, who provided Quickscat data, the SAFRAN product and ERA-interim reanalysis. The authors are finally grateful to the two anonymous reviewers who provided useful recommendations to improve the manuscript.

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