Hélène Cheap-Charpentier
email: helene.cheap-charpentier@epf.fr

Olivier Horner

Jean Lédion

Hubert Perrot

Study of the influence of the supersaturation coefficient on scaling rate using the pre-calcified surface of a quartz crystal microbalance

Keywords: calcium carbonate, scaling propensity, quartz crystal microbalance, pre-calcified surface QCM: Quartz Crystal Microbalance with a regular bare electrode surface, SQCM: Quartz Crystal Microbalance with a pre-calcified electrode surface, TOC: Total Organic Carbon, XRD: X-Ray Diffraction

Scale deposition is a common issue in industrial plants, which creates technical problems, i.e. reduction of heat transfer, decrease of flow rate due to an obstruction of pipes.

Therefore, the development of some appropriate methods based on well suitable in situ sensors to evaluate and predict the scaling propensity of water is a major concern in current research. This would be a good strategy for the optimization of anti-scaling treatments.

In this study, scaling tests were carried out using a sensitive sensor, which has been developed using a quartz crystal microbalance with a pre-calcified electrode surface (SQCM). This technique allowed studying the influence of the supersaturation on the scaling rate. The set-up was tested with different water samples which were brought to a given supersaturation coefficient by degassing the dissolved CO 2 . The prediction of the scaling propensity of water was then possible through the relationship between the scaling rate on a pre-calcified surface and the supersaturation coefficient. In addition, the kinetics of CaCO 3 deposit on the precalcified SQCM surface was found to be slower for natural water than for synthetic water (same calcium concentration). Furthermore, the activation energy for scale deposit, in synthetic water, was found to be 22 kJ.mol -1 , which may be related to the diffusion of ions and/or CaCO 3 nuclei in solution.

Introduction

Scale deposit on solid surfaces causes serious problems in many industries using or processing water, e.g. in cooling systems (Abd-El-Khaled and Abd-El-Nabey, 2013) or in heat exchangers [START_REF] Yang | Investigation of induction period and morphology of CaCO 3 fouling on heated surface[END_REF]. Indeed, the thickness of the scale deposit may limit thermal exchanges or the water flow in pipes. The use of chemicals is a common approach to control scale deposit. However, those chemicals may have an important impact on environment, i.e. in the eutrophication process. On the one hand, scale deposit could occur when water becomes supersaturated, i.e the product of ions activities, [Ca 2+ ] × [ ], exceeds the solubility product K sp of CaCO 3 [START_REF] Xyla | The precipitation of calcium carbonate in aqueous solutions[END_REF]. On the other hand, scale deposit may take place due to an external cause, i.e. the introduction of a substrate or seed crystals [START_REF] Donnet | Contribution of aggregation to the growth mechanism of seeded calcium carbonate precipitation in the presence of polyacrylic acid[END_REF]. Thus, preventing scale formation has raised a great interest. It is relevant to develop some efficient methods based on in situ sensors to follow the scaling propensity of water in order to prevent scale deposit.

Some methods based on the determination of the scaling rates have been developed in order to estimate the scaling propensity of water [START_REF] Hui | Evaluation methods for the scaling power of water[END_REF][START_REF] Leroy | Caractérisation du pouvoir entartrant des eaux à l'aide d'essai d'électrodépositionétude comparative de plusieurs méthodes[END_REF][START_REF] Al Nasser | Scaling and aggregation kinetics determination of calcium carbonate using inline technique[END_REF]. For example, Hui et al. [START_REF] Hui | Evaluation gravimétrique des vitesses d'entartrage sur tubes témoins[END_REF] have studied the scaling rates using witness tubes with pre-calcified surface. Zhang et al. (Zhang et al., 2001) considered thermodynamic models in order to evaluate the tendency for scaling from solutions and kinetics models to predict the rate of scaling. The use of chronoamperometry at constant reduction potential was the more fruitful approach to investigate the scaling rate [START_REF] Gabrielli | Estimation of the deposition rate of thermal calcareous scaling by the electrochemical impedance technique[END_REF][START_REF] Lédion | Détermination du caractère incrustant d'une eau par un essai d'entartrage accéléré[END_REF]. However, correlation between supersaturation coefficient and scaling rate on a solid surface was not very well developed in literature. Some discrepancies were detected between real deposit on solid surfaces and scaling rates estimated from predictive models based on thermodynamic data where precipitation tendency were predicted [START_REF] Hasson | Influence of the flow system on the inhibitory action of aCO 3 scale prevention additives[END_REF].

Recently, a quartz crystal microbalance with a pre-calcified sensitive surface (SQCM) has been developed (Chao et al., 2014a). The sensitive surface could be used as an adsorption layer for CaCO 3 nuclei present in water. It was installed in a cell simulating an industrial fluid stream with a laminar flow. A preliminary relationship was obtained between the instantaneous scaling rate and the supersaturation coefficient of synthetic water at different concentrations of Ca 2+ ([Ca 2+ ] in the range 60 mg.L -1 -200 mg.L -1 ) and at a given temperature (T=35°C). This opens the possibility to quantify the scaling propensity of a tested water.

In this work, the pre-calcified QCM electrode was installed in a miniaturized lab-made cell. This original set-up was tested with synthetic water which was brought to a given supersaturation coefficient related to the precipitation of calcium carbonate by CO 2 degassing [START_REF] Gauthier | Application of the Fast Controlled Precipitation method to assess the scale-forming ability of raw river waters[END_REF]Chao, 2014b). Indeed, only calcium carbonate (CaCO 3 ), one of the most abundant mineral scale compound [START_REF] Keysar | Effect of surface roughness on the morphology of calcite crystallizing on mild steel[END_REF], will be considered in this study. This work is focused on calcium carbonate scaling, i.e. the deposit of calcium carbonate on a surface, and not with calcium carbonate precipitation in solution, i.e. the appearance of a Ca CO 3 solid phase within the bulk solution. The deposit kinetics of calcium carbonate nuclei on the active electrode of the SQCM was measured at different supersaturation coefficients for two temperatures (T=30°C and 40°C), either for a synthetic or a natural water. Indeed, it was possible to compare, in a very sensitive way, the scaling propensity of a synthetic and a natural mineral water (same calcium concentration). Thus, this work investigates, thanks to the pre-calcified QCM electrode, the effect of several parameters like supersaturation level, temperature and composition of water on scaling rate. For the first time, the activation energy of the scaling process on the pre-calcified surface, for synthetic water, was investigated.

Materials and methods

Reactants

Synthetic waters were pure solutions, which only contained Ca 2+ and CO

QCM set-up

A quartz crystal microbalance with a pre-calcified electrode surface (SQCM) was used to measure the mass of calcium carbonate deposited on the pre-calcified surface (Figure 1). The volume of the SQCM cell was 500 L. The flow rate was set to 2 mL.min -1 using a peristaltic pump. The water jet arrived at 45° on the flat surface of the modified electrode.

The bare surface of the electrode consisted in a 5 mm diameter gold disc covering a 9 MHz AT cut quartz crystal resonator (AWS, Spain).

The pre-calcifying process was performed in a classical "three-electrode" set-up with the bare gold electrode acting as a working electrode. All the electrodes were inserted in a submerged impinging jet cell (flow rate of 300 mL.min -1 ), with a synthetic water (200 mg.L -1

Ca 2+ ) at 30°C.

The frequency of the microbalance device was measured by a frequency-counter (Fluke PM6685). The mass deposited onto the bare or the pre-calcified surface of the sensor was measured over time from the microbalance frequency change [START_REF] Bizet | Immunodetection by quartz crystal microbalance -A new approach for direct detection of rabbit IgG and peroxidase[END_REF]Chao et al., 2014a). The frequency change f of the quartz resonator due to the deposition of the scale on the working electrode was proportional to the mass change m according to Sauerbrey equation [START_REF] Sauerbrey | Verwendung von Schwingquarzen zur Wägung dünner Schichten und zur Mikrowägung[END_REF]:

f = -K s .m (1)
where K s is the sensitivity factor (K s ) -1 = 1,09 ng per Hz for an electrode area of 0.2 cm 2 .

Preparation of water with a given supersaturation coefficient 

Stable solutions of calcium carbonate were brought to a given supersaturation coefficient ( coefficient) by CO 2 degassing (Lédion et al., 1997;[START_REF] Gauthier | Application of the Fast Controlled Precipitation method to assess the scale-forming ability of raw river waters[END_REF]Chao et al., 2014b). Indeed, the degassing of dissolved CO 2 is accelerated by stirring the water sample. As the CO 2 degasses, the concentration of [H + ] decreases, so the concentration of [OH -] or the pH increases according to the following equations:

CO 2 dissolved ↔ CO 2 atmospheric (2) HCO 3 -↔ CO 2 dissolved + OH - (3) H + + OH -↔ H 2 O (4)
As a consequence, the saturation level of the solution increases with stirring.

Supersaturation coefficient , the extent of deviation of the solution from the equilibrium condition, is the driving force for the formation of a CaCO 3 crystalline phase in aqueous solutions.  is defined according to the following equation ( 5): [START_REF] Legrand | Les équilibres carboniques et l'équilibre calcocarbonique dans les eaux naturelles[END_REF])

 = [Ca 2+ ]  [CO 3 2- ] / K' sp (5)
where [Ca 2+ ] and [CO 3

2-

] are the concentrations of calcium ions and carbonate ions respectively, and K' sp is the apparent solubility product of calcium carbonate. As stated above, the supersaturation level  of the studied water was increased by using a moderate stirring (400 rpm) of the solution. For each test, the synthetic water was initially brought to the same supersaturation coefficient (=40), below the homogeneous precipitation threshold. The pH of the solution was measured as a function of time using a pH-meter (Radiometer pHM220). The pH electrode was purchased from Radiometer Analytical, and the saturated calomel electrode (Radiometer Analytical) was used as a reference electrode. When the solution reached a given  value (i.e. a given pH value), the stirring was stopped and all the openings of the beaker were hermetically sealed. The value of  was maintained below the spontaneous precipitation threshold.

X-ray diffraction (XRD)

The XRD method was used in order to determine crystalline polymorphs of CaCO 3 formed on the gold surface. XRD measurements were performed at room temperature with Cu-Kα radiation (1.52 Å) using a Panalytical (Empyrean Diffractometer Panalytical) device with prefix configuration. Crystalline forms of the CaCO 3 layer forming the pre-calcified surface were determined from spectra obtained by XRD measurements.

Scaling rate equation

During this scaling process, the balance between precipitation and dissolution of scaling can be described by the following equation:

k k R P     (6)
where R represents Ca 2+ and CO 3 2-species and P represents calcium carbonate in solid phase.

k + and k -are the rate constants related to the reactions of precipitation or dissolution, respectively. The relationship between the reaction rate, V S , and the free energy, G is given by equation ( 7): [START_REF] Blum | The effect of dislocation density on the dissolution rate of quartz[END_REF]

where V S is the reaction rate, S is the reaction surface, V is the volume of the solution, ΔG pre and ΔG dis are, respectively, the free energy released in CaCO 3 formation process and the free energy expended in dissolution of calcium carbonate, ΔG is the total free energy of the reaction, k is the rate constant of the reaction and n the reaction order. G depends on the supersaturation level, δ, of the solution according to the following equation:

2 g RT G ln   (8)
For the scaling rate (δ>1), equation 7 can be rewritten as:

log Vs = log k' + n  log [ 1/2 -1] (9)
where V S (µg.cm -2 .min -1 ) is the scaling rate, k' is the rate constant for deposit (k' = (S/V) k)

and n is the apparent reaction order [START_REF] Morse | A calcium carbonate formation and dissolution[END_REF].

Results and discussion

Pre-calcification of the sensitive surface of QCM

The sensitive surface of QCM can be initially covered by a calcium carbonate layer to obtain the SQCM sensor [START_REF] Gabrielli | Quartz crystal microbalance investigation of electrochemical calcium carbonate scaling[END_REF][START_REF] Peronno | Study of the inhibition effect of two polymers on calcium carbonate formation by fast controlled precipitation method and quartz crystal microbalance[END_REF]. The potential applied to the working electrode was -1 V/SCE, leading to the reduction of the dissolved dioxygen to produce hydroxyl ions. The formed hydroxyl ions increased the local pH near the electrode surface [START_REF] Tlili | Influence of the interfacial pH on electrochemical CaCO 3 precipitation[END_REF] which led to the transformation of bicarbonate ions into carbonate ions. The carbonate ions reacted with calcium ions to form a CaCO 3 precipitate on the surface of the working electrode.

The pre-calcifying experiment was stopped when the current, recorded during calcium carbonate deposition, reached a low residual constant value of a few micro amps. This indicates that the electrode surface was fully recovered by the CaCO 3 film. In these experimental conditions, the solid phase formed on the electrode surface was pure calcite, the most thermodynamics form of calcium carbonate [START_REF] Morse | A calcium carbonate formation and dissolution[END_REF], as shown in Figure 2: Obtaining pure calcite is likely to be due to a moderate kinetics of deposition. If the kinetics of deposition would increase, a mixture of calcite and aragonite would be observed.

This process of layer formation is highly reproducible and the CaCO 3 layer was stable over time [START_REF] Gabrielli | Nucleation and growth of calcium carbonate by an electrochemical scaling process[END_REF]. In Figure 3b, the mass change detected on the bare gold surface was roughly constant.

Comparison of the

This indicates that no CaCO 3 deposition was detected by the QCM sensor with no precalcified surface. On the contrary, CaCO 3 deposit is immediately observed on the precalcified surface of the sensor (Figure 3a). It should be noticed that the mass deposited on the pre-calcified surface increased linearly over time.

In this work, the pre-calcified surface of the sensor is very sensitive towards CaCO 3 nuclei adsorption, compared to a bare gold surface. This result validated the use of a precalcified surface for scaling tests. This is in line with the results obtained from the literature. [START_REF] Hui | Evaluation gravimétrique des vitesses d'entartrage sur tubes témoins[END_REF] on titanium or copper tubes. Indeed, they showed that the scaling rate was enhanced when the inner surface of the tubes was pre-calcified. Figure 3 shows that the mass change on the pre-calcified surface (pure calcite) depends strongly on the supersaturation level of the studied solution (with the others parameters constant). The higher was the degree of supersaturation, the faster were the scaling rate and the amounts of deposit. This suggests that the number of CaCO 3 nuclei strongly affects the scaling rates (Chao et al., 2014a). Moreover, SQCM could detect CaCO 3 nuclei, even at very low supersaturation coefficient (=4). Recently, Chao et al. showed, by using Small Angle X-rays Scattering (SAXS) spectroscopy, that CaCO 3 nuclei were present in synthetic water before the homogeneous precipitation (Chao et al., 2014b). These CaCO 3 nuclei can be detected also in the early stage of the experiments with this measurement (Demichelis et al., 2011). It should be noticed that the scaling rate V s , on a pre-calcified surface, was constant over time for given experimental conditions. For each  value, the scaling rate V s can be calculated from the slope of the linear regression related to the mass change vs time curve (Fig. 3). V s characterizes the scaling propensity of water (Chao et al., 2014a). A plot of the logarithm of scaling rate as a function of the logarithm of the supersaturation coefficient could be linearly fitted using equation 9, as detailed in the Materials and Methods part (Figure 5). (▲) 40°C (flow rate=2 mL.min -1 , pre-calcified surface of the sensor made of pure calcite).

Supersaturation level vs scaling rate for synthetic water

The data related to the linear fit are summarized in Table 2. As shown in Figure 4, the scaling rates of calcium carbonate were found to depend strongly on the value of the supersaturation level of the studied solution, for a given water (Zhang et al., 2001;[START_REF] Chen | Calcium carbonate scale formationassessing the initial stages of precipitation and deposition[END_REF][START_REF] Parlaktuna | The use of chemical inhibitors for prevention of calcium carbonate scaling[END_REF][START_REF] Abdel-Aal | Study of the adhesion mechanism of CaCO 3 using a combined bulk chemistry/QCM technique[END_REF][START_REF] Ben Amor | Influence of water hardness, substrate nature and temperature on heterogeneous calcium carbonate nucleation[END_REF]. The scaling rate increased with in agreement with the results from literature. For example, Al Nasser et al. [START_REF] Al Nasser | Scaling and aggregation kinetics determination of calcium carbonate using inline technique[END_REF] showed that the deposit rate on steel plates immerged in a scale forming solution increased when  increased. Koutsoukos and Kontoyannis [START_REF] Koutsoukos | Precipitation of calcium carbonate in aqueous solutions[END_REF] measured the scaling rate of calcium carbonate precipitation from aqueous solution by following the disappearance of calcium ion over time. They showed that the scaling rate increased with the supersaturation coefficient.

It must be noted that the orders of reaction n shown in Table 2 are of the same order of magnitude as those reported in literature. Indeed, an apparent order of 1.3 (T = 22°C) was reported for the CaCO 3 growth on nuclei [START_REF] Packter | The precipitation of sparingly soluble alkaline-earth metal and lead salts: nucleation and growth orders during the induction time[END_REF]. Moreover, [START_REF] Kitano | The influence of organic material on the polymorphic crystallization of calcium carbonate[END_REF] found a reaction order of ca 1 for the growth of CaCO 3 . Parsiegla and Katz [START_REF] Parsiegla | Calcite growth inhibition by copper(II). I. Effect of supersaturation[END_REF] have reported an apparent order of 1 for calcite growth at low pH and low carbonate concentration. However, in their experiment they have investigated the kinetics of calcite growth. In this work, the CaCO 3 layer was immobilized on a gold surface, and the scaling rate is related to the adsorption of nuclei formed in solution, likely as in a real scaling phenomena. In "natural conditions", scaling phenomena occur on a surface for moderate supersaturation levels.

Influence of the temperature on scaling rate in synthetic water

SQCM was performed at higher temperature. The choice of the higher but moderate temperature (40°C) prevent the risk of precipitation of water at the beginning of the experiment. Figure 4 shows that the scaling rates established at 40°C were higher than those obtained at 30°C. This result is in agreement with previous studies (Al [START_REF] Al Nasser | Scaling and aggregation kinetics determination of calcium carbonate using inline technique[END_REF][START_REF] Ben Amor | Influence of water hardness, substrate nature and temperature on heterogeneous calcium carbonate nucleation[END_REF]. Indeed, the solubility of CaCO 3 decreases when the temperature increases [START_REF] Plummer | The solubilities of calcite, aragonite and vaterite in CO 2 -H 2 O solutions between 0 and 90°C, and an evaluation of the aqueous model for the system CaCO 3 -CO 2 -H 2 O[END_REF] and the temperature dependence on the scaling rate may affect the growing process on the pre-calcified surface [START_REF] Hillig | A derivation of classical two-dimensional nucleation kinetics and the associated crystal growth laws[END_REF].

According to the Arrhenius law, the calculated activation energy (from temperatures T=30°C and 40°C) in the experimental conditions of this work was found to be 22 kJ.mol -1 .

In literature, values of activation energies have been reported for surface controlled processes involving CaCO 3 precipitation on crystal seeds. For example, Koutsoukos and Kontoyannis [START_REF] Koutsoukos | Precipitation of calcium carbonate in aqueous solutions[END_REF] found an apparent activation energy of 155 kJ.mol -1 for a growth process on CaCO 3 seeds, suggesting that the calcium carbonate precipitation was surface-controlled in aqueous solution. An activation energy of 46 kJ.mol -1 was also reported for crystal growth on seed crystals of calcite [START_REF] Nancollas | The crystallization of calcium carbonate, II. Calcite growth mechanism[END_REF][START_REF] Cassford | Crystallisation kinetics of calcite from calcium bicarbonate solutions between 278[END_REF][START_REF] Xyla | The precipitation of calcium carbonate in aqueous solutions[END_REF]. In the present work, it must be noticed that the activation energy (22 kJ.mol -1 ) is related to a scaling process on an immobilized CaCO 3 layer.

The influence of the flow rate on scale deposition was further investigated, in order to see if the activation energy determined in this study is related to diffusion (Fig. S2 in Supplementary Information). It was shown, by using the SQCM set-up with synthetic water (T=30°C, =20, pre-calcified surface of the sensor made of pure calcite) that the kinetics of scale deposition increased with the flow rate. This suggests that the kinetics of scale deposition may be controlled by the diffusion of ions and/or CaCO 3 nuclei in solution. In a previous work, the activation energy related to the dissolution process of vaterite was found to be 24 kJ.mol -1 and was typical of the ion diffusion in aqueous solution [START_REF] Brecevic | On calcium carbonates: from fundamental research to application[END_REF].

Natural water vs synthetic water

The Salvetat water is a natural water, which is very close to synthetic water, i.e. a comparable amount of bicarbonate and calcium, and a very low content in other ions, especially Mg 2+ and Cl -ions. The FT-IR spectrum of the precipitate obtained from the Salvetat water is very similar to that of the pure calcium carbonate used for synthetic water (Fig. S1 in Supplementary Information). In addition, there are no silicates or phosphates species detectable in Salvetat water precipitate. As the Salvetat water is naturally aggressive, the presence of CaCO 3 germs at the beginning of the SQCM experiment is avoided. Thus, the Salvetat water can be used to show the influence on scaling rate of elements that are not well characterized but potentially present in water, i.e. natural organic substances and/or suspended matter.

The mass change of the scale quantity deposited on the pre-calcified surface from a synthetic water ([Ca 2+ ]=200 mg.L -1 ) and from a natural water (Salvetat, [Ca 2+ ]=180 mg.L -1 ) were investigated by using the SQCM set-up (Figure 6):

A B

Figure 6A shows that the SQCM is sensitive enough to determine the mass change in natural water. The results showed that the scaling rate is lower for the natural water (curve b) than for the synthetic water (curve a). A previous study using electrochemical QCM (EQCM) has shown that organic molecules with carboxylic groups were very efficient to inhibit towards calcium carbonate scaling [START_REF] Peronno | Study of the inhibition effect of two polymers on calcium carbonate formation by fast controlled precipitation method and quartz crystal microbalance[END_REF]. The effect of humic substances (Chao et al., 2014a) and kaolinite [START_REF] Chao | Scaling Potential determination of water circuit by the use of an ultrasensitive crystal quartz microbalance[END_REF] were investigated by using SQCM. It was shown that the average scaling rate decreased when the humic substance concentration increased. SQCM also showed that kaolinite had little influence on the scaling propensity of water. A slowdown of nucleation kinetics was also observed in the presence of clay suspensions [START_REF] Hui | Influence of clay suspensions on scaling[END_REF].

As previously discussed the Salvetat water is, very close to synthetic water but also contains some elements that are not well characterized as organic substances (TOC of 0.4 mg/L) or inorganic substances. The data shown in Figure 6 shows that those additional elements which decrease the scaling rate. These elements could corresponds to organic molecules, humic substances or suspended matter (clay...), in line with the results of the literature [START_REF] Ladel | Mise en évidence de l'effet inhibiteur de métabolites d'algues planctoniques sur la précipitation du carbonate de calcium dans les eaux naturelles d'origine superficielle[END_REF]. They may poison the active sites of the CaCO 3 crystals and thus strongly influence the deposition process. The surface charges of these elements could play a role in those mechanisms. For example, Hui et al. showed that the antiscaling effect of clays is maximum if the clay displays the maximum quantity of suspended materials as well as a highly negative zeta (ζ) potential [START_REF] Hui | Influence of clay suspensions on scaling[END_REF].

Figure 6B shows a comparison of the mass change of scale deposition for Salvetat water (curve a) and Salvetat water diluted at 50% (curve b). The supersaturation coefficient is slightly higher for Salvetat water diluted at 50% than for Salvetat water (9.60 and 8.45 respectively). It must be noticed that the scaling rate is higher for Salvetat water than for Salvetat water diluted at 50%. Indeed, the Ca 2+ concentration is lower, and the scaling rate decreased when the concentration of calcium decreased, i.e. the number of CaCO 3 nuclei decreased. The supersaturation coefficient is therefore not the only one parameter which influences the scaling rate.

Conclusion

In this work, a scaling quartz crystal microbalance (SQCM) with a pre-calcified surface constituted of pure calcite was used to assess the scaling propensity of synthetic or natural water.

The scaling rate on the pre-calcified surface was markedly dependent on solution supersaturation coefficient and temperature; it increased with the increase of these two parameters. For the first time, the activation energy related to a scaling process on an immobilized CaCO 3 layer (22 kJ.mol -1 ) was evaluated by using SQCM. The kinetics of scale deposition increased with the flow rate. This phenomenon may be therefore controlled by the diffusion of ions and/or CaCO 3 nuclei in solution.

Finally, the SQCM set-up is sensitive enough to show that scaling rate is lower in natural water, compared to synthetic water in similar conditions. This may be due to the presence of the organic compounds or particle suspensions in natural water which could have some significant inhibition effects.

This work confirms that SQCM is a convenient and sensitive tool to measure in situ CaCO 3 nuclei deposit both in synthetic and natural water. Even at the beginning of the experiment with small CaCO 3 particles at low concentration, it could be used as an in situ sensor to assess the scaling propensity for a given water, immerged in a main circuit or in a bypass. This could be interesting for some industrial applications (i.e. cooling system or circuit). It could also be used profitably to assess efficiency of an antiscalant treatment by considering the scaling propensity of water.

Figure 1 :

 1 Figure 1: Scheme of a SQCM set up.

Figure 2 :

 2 Figure 2: XRD spectrum of the CaCO 3 layer formed during the pre-calcifying process.

CFigure 3 :

 3 Figure 3 shows the curves of mass change as a function of time for synthetic water ([Ca 2+ ]=200 mg/L, T=30°C) for a pre-calcified surface (pure calcite), and a bare gold surface

  Kinetics of calcium carbonate nuclei deposit on a pre-calcified surface was determined by the SQCM device for synthetic water with given  coefficients (see Materials and methods). Figure4shows the influence of the supersaturation level over the mass change on a pre-calcified surface (pure calcite) as a function of time.

Figure 4 :

 4 Figure 4: Mass change as a function of time: influence of the supersaturation coefficient (T=30°C, flow rate=2 mL.min -1 , pre-calcified surface of the sensor made of pure calcite).

Figure 5 :

 5 Figure 5: Logarithm of the scaling rate V S as a function of the logarithm of  at (■) 30°C and

Figure 6 :

 6 Figure 6: (A) Mass change of scale deposited from (a) synthetic water ([Ca 2+ ]=200 mg.L -1 ,

Figure captions Figure 1 :

 captions1 Figure captions

Figure 2 :

 2 Figure 2: XRD spectrum of the CaCO 3 layer formed during the pre-calcifying process. C:

Figure 3 :

 3 Figure 3: Curves of mass change as a function of time for synthetic water with (a) a precalcified surface (pure calcite) (scaling time Vs=1.2 µg.cm -2 .min -1 ) and (b) a bare gold surface (initial [Ca 2+ ]=200 mg/L, T=30°C, flow rate=2 mL.min -1 , =40).

Figure 4 :

 4 Figure 4: Mass change as a function of time: influence of the supersaturation coefficient (T=30°C, flow rate=2 mL.min -1 , pre-calcified surface of the sensor made of pure calcite).

Figure 5 :

 5 Figure 5: Logarithm of the scaling rate V S as a function of the logarithm of  at (■) 30°C and(▲) 40°C (flow rate=2 mL.min -1 , pre-calcified surface of the sensor made of pure calcite).

Figure 6 :

 6 Figure 6: (A) Mass change of scale deposited from (a) synthetic water ([Ca 2+ ]=200 mg.L -1 ,

Table 1 :

 1 Concentration of ions in Salvetat® water.

	3	2-ions. The

Table 2 :

 2 Summary of data obtained from fitted curves (see text for signification of the

	parameters).
	T (°C)	k/SI	n	R 2*
	30	0.19	1.3	0.95
	40	0.36	1	0.98

* Goodness-of-fit measure for the linear regression model.

Table captions Table 1 :

 captions1 Concentration of ions in Salvetat® water.

Table 2 :

 2 Summary of data obtained from fitted curves (see text for signification of the parameters).

Acknowledgement

Cyrille Bazin (LISE) is gratefully thanked for technical assistance on X-Ray diffraction experiment. Axel Desnoyers de Marbaix (LISE) is thanked for the manufacturing of the QCM cell.

Zhang, Y., Shaw, H., Farquhar, R., Dawe, R., 2001. The kinetics of carbonate scaling -1 application for the prediction of downhole carbonate scaling. J. Petrol. Sci. Eng. 29,[85][86][87][88][89][90][91][92][93][94][95]


Adolfo Chamorro
email: achamorro@imarpe.gob.pe

Vincent Echevin

Francois Colas

Vera Oerder

Jorge Tam

Carlos Quispe-Ccalluari

Mechanisms of the intensification of the upwelling-favorable winds during El Niño 1997-1998 in the Peruvian Upwelling System

Keywords: Ocean-atmosphere interactions, Coastal winds, El Niño 1997/1998, Peruvian Upwelling System

The physical processes driving the wind intensification in a coastal band of ~100 km off Peru during the intense 1997-1998 El Niño (EN) event were studied using a regional atmospheric model. A simulation performed for the period 1994-2000 reproduced the coastal wind response to local sea surface temperature (SST) forcing and large scale atmospheric conditions. The model, evaluated with satellite data, represented well the intensity, seasonal and interannual variability of alongshore (i.e. NW-SE) winds. An alongshore momentum budget showed that the pressure gradient was the dominant force driving the surface wind acceleration. The pressure gradient tended to accelerate the coastal wind, while turbulent vertical mixing decelerated it. A quasilinear relation between surface wind and pressure gradient anomalies was found. Alongshore pressure gradient anomalies were caused by a greater increase in near-surface air temperature off the northern coast than off the southern coast, associated with the inhomogeneous SST warming. Vertical profiles of wind, mixing coefficient, and momentum trends showed that the surface wind intensification was not caused by the increase of turbulence in the planetary boundary layer. Moreover, the temperature inversion in the vertical mitigated the development of pressure gradient due to air convection during part of the event. Sensitivity experiments allowed to isolate the respective impacts of the local SST forcing and large scale condition on the coastal wind intensification. It was primarily driven by the local SST forcing whereas large scale variability associated with the South Pacific Anticyclone modulated its effects. Examination of other EN events using reanalysis data confirmed that intensifications of alongshore wind off Peru were associated with SST alongshore gradient anomalies, as during the 1997-1998 event.

Introduction

The Peruvian Upwelling System is one of the major upwelling systems of the world in terms of fisheries (Zuta and Guillen 1970;[START_REF] Chavez | The northern Humboldt Current System: Brief history, present status and a view towards the future[END_REF]. A key oceanic process in this nearshore marine environment is the upwelling of deep, nutrient-replete, cold water to the surface forced by Ekman divergence associated with predominantly equatorward coastal winds. As in other upwelling regions, it is characterized by an intense biological productivity [START_REF] Chavez | A comparison of eastern boundary upwelling ecosystems[END_REF] and strong air-sea interactions (e.g. Halpern 2002;[START_REF] Boé | What shapes mesoscale wind anomalies in coastal upwelling zones?[END_REF][START_REF] Oerder | Mesoscale SSTwind stress coupling in the Peru-Chile current system: Which mechanisms drive its seasonal variability?[END_REF]. A unique aspect of the Peruvian system is its proximity to the equator in the Eastern Pacific, which places it directly under the influence of El Niño events (EN hereafter). During the so called "canonical" EN events, warm surface waters accumulate in the Eastern Tropical Pacific off the Ecuador and Peru coasts (e.g. [START_REF] Picaut | Mechanisms of the 1997-1998 El Niño-La Niña, as inferred from space-based observations[END_REF] causing a dramatic reduction of the upwelling of cold water (e.g. [START_REF] Colas | 1997-98 El Niño off Peru: A numerical study[END_REF]. The upwelling reduction is somewhat mitigated by an increase of the equatorward coastal wind (Wyrtki 1975;[START_REF] Kessler | The circulation of the eastern tropical Pacific: A review[END_REF]). This local wind increase is seemingly paradoxical since large-scale trade winds are weakened in the equatorial [START_REF] Bjerknes | Survey of El Niño 1957-58 in its relation to tropical Pacific meteorology[END_REF]) and subtropical [START_REF] Rahn | Influence of large scale oscillations on upwelling-favorable coastal wind off central Chile[END_REF] regions. Figure 1a shows the mean spatial distribution of the wind anomalies off Peru coast during the strongest El Niño event observed to date, between November 1997 and February 1998. Positive wind anomalies were maximum onshore and decreased offshore. Negative wind anomalies at the equator indicate the weakening of the southerly trade winds in the equatorial Pacific. The strongest alongshore positive wind anomalies reached ~1.5 m s -1 during November 1997-February 1998 which represents a ~40 % increase with respect to the mean climatological conditions (Figure 1b). Note that in the present work, anomalies were computed with respect to mean climatological conditions over the 1994-2000 period.

The processes that drive the nearshore wind increase during EN have not been studied in detail. Previous studies suggested that the wind intensification could be driven by a strengthening of the cross-shore pressure gradient (supporting a geostrophic wind) owing to an enhanced cross-shore thermal contrast between land and sea (Bakun 1990;[START_REF] Bakun | Greenhouse gas, upwelling favourable winds, and the future of upwelling systems[END_REF]. This enhanced thermal contrast would be caused by a stronger temperature increase over land than over sea during EN, due to the greenhouse effect induced by moist air. On the other hand, Enfield (1981) suggested that the enhanced cross-shore thermal contrast during EN was forced by a stronger shortwave heating over land associated with a reduction of nearshore cloudiness. However, SST gradients may also impact on surface winds. [START_REF] Lindzen | On the role of sea surface tem-perature gradients in forcing low level winds and convergence in the Tropics[END_REF] showed that surface winds over the tropical Pacific can be forced by SST gradients at relatively large scale. [START_REF] Kessler | The circulation of the eastern tropical Pacific: A review[END_REF] suggested that the alongshore SST gradient which appears off the Peru coasts during EN may drive a strengthening of the alongshore pressure gradient and wind, but the mechanisms were not studied. In addition, the enhanced atmospheric turbulence due to the surface ocean warming may result in the downward vertical mixing of momentum from the upper layers of the atmosphere to the ocean surface, thus increasing the surface wind (e.g. [START_REF] Wallace | The influence of sea surface temperature on surface wind in the Eastern[END_REF].

Large scale atmospheric circulation is also impacted by EN, which may affect coastal winds through modifications of the South Pacific Anticyclone (SPA). Dewitte et al. (2011) showed that the alongshore wind intraseasonal (i.e. 10-60 days band) variability off central Peru (~15°S) was forced by migratory disturbances across the SPA. [START_REF] Rahn | Influence of large scale oscillations on upwelling-favorable coastal wind off central Chile[END_REF] showed that the SPA was weaker during EN, resulting in decreasing winds off central Chile. Such weakening might mitigate the coastal wind increase during EN off Peru, and a poleward displacement of the SPA might have a similar effect, as shown by [START_REF] Belmadani | What dynamics drive future wind scenarios for coastal upwelling off Peru and Chile?[END_REF] in the context of climate change. However, the influence of large-scale atmospheric circulation interannual variability on the coastal winds off Peru remains to be extensively investigated.

In this paper, a regional atmospheric model forced by realistic oceanic (i.e. Sea Surface Temperature) and lateral boundary conditions was used to investigate the physical processes driving the coastal wind intensification during the strong 1997-1998, "eastern pacific" El Niño event. The respective roles of the large scale signal and SST local forcing on the alongshore wind anomalies were also studied for this particular event. Data and methods are described in section 2. Results are presented in section 3, and discussion and conclusions are given in section 4.

Data and Methods

Regional atmospheric model

The Weather Research and Forecasting (WRF) model version 3.3.1 [START_REF] Skamarock | A time-split nonhydrostatic atmos-pheric model for weather research and forecasting applications[END_REF] was used to simulate the coastal wind during the period [START_REF] Janjic | The step-mountain eta coordinate model: further developments of the convection, viscous sublayer, and turbulence closure schemes[END_REF]-2000, which includes the very strong 1997[START_REF] Strub | The global coastal ocean: Regional studies and syntheses[END_REF]El Niño event (McPhaden 1999). WRF is a fully compressible and nonhydrostatic model. Its vertical coordinate is a terrainfollowing hydrostatic pressure coordinate. The model uses a time-split integration scheme. Slow and lowfrecuency modes are integrated using a Runge-Kutta 3rd order time integration scheme, while the high-frecuency acoustic modes are integrated over smaller time steps to maintain numerical stability. For spatial discretization the model uses a 5th order upwind biased advection schemes. Two nested domains, with one-way offline nesting, were used (Fig. 2). The large domain has a resolution of 0.75° and encompasses the South East Pacific and the main part of South America. The small domain has a resolution of 0.25° and covers the Peru and Northern Chile region (30°S-12°N). Both domains include the Andes. The relatively high resolution of the nested domain allows to better represent the orography (Fig. 2), a crucial element of the regional dynamics (e.g. Xue et al. 2004). Both grids have 60 vertical sigma levels between the surface and the top of the atmosphere (defined by the 50 hPa top pressure), with 21 levels in the first ~1000 m. The time steps for the large and nested domains are 180 s and 60 s, respectively.

The parameterizations for short and long wave radiation, cloud physics, land surface and planetary boundary layer (PBL) used in this study are listed in Table 1. Most of them (except the [START_REF] Dudhia | Numerical study of convection observed during the winter monsoon experiment using a mesoscale two-dimensional model[END_REF] shortwave radiation scheme) are identical to those in the 1/12° configuration of [START_REF] Oerder | Mesoscale SSTwind stress coupling in the Peru-Chile current system: Which mechanisms drive its seasonal variability?[END_REF] for the Peru-northern Chile region. The ERA-interim reanalysis data [START_REF] Dee | The ERA-Interim reanalysis: configuration and performance of the data assimilation system[END_REF]), 6-hourly, were used as initial and boundary conditions. The daily Optimum Interpolation Sea Surface Temperature (OISST) at 0.25° [START_REF] Reynolds | Daily high-resolution-blended analyses for sea surface temperature[END_REF] was used as SST forcing. Diurnal SST variations are accounted for in our simulations by using the Zeng and Beljaars (2005) slab model that is included in WRF. Model outputs were recorded every six hours using instantaneous and average values.

Observational data

Observations from two different satellite-borne scatterometers were used to evaluate the realism of the model surface winds: the ERS weekly wind fields at 1°x1° resolution over the period 1992-2000 and a monthly climatology of QuikSCAT wind fields (hereafter QSCAT) at 1/2°x1/2° resolution grids. The two products were processed by CERSAT (2002a,b) and downloaded from www.ifremer.fr/cersat. Surface winds were interpolated on the 1/4° model grid.

Daily SST and output from ERA-interim reanalysis (see above) were also used in the analysis.

Monthly momentum budget

In order to investigate the dominant forces that induce the monthly wind changes, we used the following relation demonstrated in [START_REF] Madec | NEMO ocean engine. Note du Pole de modélisation[END_REF] for the NEMO ocean model) and [START_REF] Oerder | Mesoscale SSTwind stress coupling in the Peru-Chile current system: Which mechanisms drive its seasonal variability?[END_REF]:

⟨𝑉⟩-𝑉(𝑡 0 ) Δt = ∑ [𝐹 𝑛 ] 𝐹 𝑛 ∈{forces} ( 1
)
where ⟨𝑉⟩ is the monthly mean wind, 𝑉(𝑡 𝑜 )is the initial velocity at the beginning of the month, Δtthe

[𝐹 𝑛 ] = 1 N+1 ∑ (∑ 𝐹 𝑛 𝑝 k=1 ) 𝑁 p=0 (2)
with N the number of time steps during one month.

Based on Eq. ( 1), we obtain the following relation for two consecutive months (M and M+1):

⟨𝑉⟩ M+1 -⟨𝑉⟩ 𝑀 =V(𝑡 0,M+1 ) -𝑉(𝑡 0,M ) + ∑ Δt 𝐹 𝑛 ∈{forces} ([𝐹 𝑛 ] M+1 -[𝐹 𝑛 ] 𝑀 ) (3)
On the other hand, a simple integration of the momentum equation between t0,M and t0,M+1 (i.e. the respective dates corresponding to the beginning of months M and M+1) leads to:

𝑉(𝑡 0,M+1 ) -𝑉(𝑡 0,M )=Δt ∑ ⟨𝐹 𝑛 ⟩ 𝐹 𝑛 ∈{Forces} (4)
Thus ⟨𝐹 𝑛 ⟩ is the time average of the forces between the beginning of the month (t0,M) and the beginning of the following month (t0,M+1).

Finally, replacing Eq. ( 4) in Eq. ( 3), we obtain:

⟨𝑉⟩ M+1 -⟨𝑉⟩ 𝑀 = ∑ Δt 𝐹 𝑛 ∈{forces} ([𝐹 𝑛 ] M+1 -[𝐹 𝑛 ] 𝑀 + ⟨𝐹 𝑛 ⟩) (5)
The left side of Eq. ( 5) represents the change of the monthly mean wind (from month M to month M+1) and the right side represents the sum of the forces contribution. We computed the anomalies of the alongshore (i.e.

parallel to the WRF model smoothed coastline oriented in the NW-SE direction) component of Eq. ( 5) (with respect to the "climatology" over 1994-2000, i.e. the average over all years of the difference of one month from its predecessor) in a coastal band of one degree width (4 grid points of the nested domain). This coastal band fully covers the upwelling area which offshore extension is controlled by the shelf topography [START_REF] Estrade | Cross-shelf structure of coastal upwelling: A twodimensional expansion of Ekman's theory and a mechanism for innershelf upwelling shut down[END_REF] and does not exceed the Rossby radius of deformation (around 120 km off Peru, e.g. [START_REF] Chavez | An estimate of new production in the equatorial Pacific[END_REF]. In this area were observed the maximum wind anomalies during the 1997-98 EN (Fig. 1a). WRF code modifications were necessary to save online the individual tendency terms of the momentum balance at each model time step. However modification of the tendency terms due to high-frequency acoustic modes were not saved, which explains why an exact closure of the momentum balance can not be expected.

Virtual temperature

Virtual temperature is computed in order to estimate the relative contribution of humidity to air density and pressure during EN. The virtual temperature of moist air is the temperature that dry air should have to reach a total pressure and density equal to those of the moist air [START_REF] Wallace | Atmospheric Science: A Introductory Survey[END_REF]. It is defined by:

𝑇 𝑣 =T * (1 + 0.61 * 𝑄 𝑣 ) (6)
with T the air temperature and Qv the mixing ratio, which describes humidity.

We computed the monthly virtual temperature anomaly as:

𝑇 𝑣 '=(1 + 0.61 * 𝑄 𝑣 ) * T'+(𝑇 * 0.61) * 𝑄 𝑣 ′ (7)

with primes marking the monthly anomalies and overlines marking the monthly means. The two terms on the right side of Eq. ( 7) represent the relative contribution of air temperature and humidity anomalies to the virtual temperature anomaly. Given that virtual temperature is directly proportional to the pressure (for a fixed mass of gas, at a constant volume), these terms contribute to the atmospheric pressure.

Model simulations

We carried out three experiments with the WRF model. First, a control run was performed over the 1994-2000 period. Then, two experiments (BRY-EN and SST-EN ) were performed to isolate the role of the large scale signal from the role of the SST local forcing. The BRY-EN experiment was carried out using the 6-hourly atmospheric boundary conditions from the years 1997-1998 (Niño boundary conditions) and daily SST forcing from the years 1994-1995 (so-called SST neutral conditions) to isolate the role of the large scale signal. The SST-EN experiment was performed using the 6-hourly atmospheric boundary conditions from the years 1994-1995 (neutral boundary conditions) and daily SST forcing from the years 1997-1998 (SST Niño conditions) to isolate the role of the SST local forcing. Note that the 1994-1995 period is considered here as "neutral conditions" for the Peru region as it did not present strong anomalies with respect to climatological conditions, in spite of the occurrence of a weak Central Pacific El Niño in austral spring 1994 and summer 1995.

Statistical significance

For each of the correlations between two variables presented in the following, the Pearson correlation coefficient (r) was used to measure the strength of a linear association between the variables, and the p-value was used to determine the statistical significance of the relationship.

Results

Validation

Mean and annual cycle of the surface wind

The simulated surface winds were compared with ERS and QSCAT observations. Figures 3 a-c show the annual mean surface wind field obtained from both satellites and the model for the year 2000, when the two satellite observation periods were overlapping. South of the equatorial line, the observed surface winds were strong (~6-8 m s⁻ ¹) and they blowed north-westward over the oceanic region. They were weaker and approximately parallel to the coastline in the nearshore region, with a maximum (~5-6 m s -1 ) near 15°S and minimum (~3-4 m s -1 ) near 18°S (Figs. 3a,c). The model reproduced the observed wind spatial patterns, with a wind intensity closest to that of QSCAT. However, the wind drop-off (i.e. wind decrease towards the coast, e.g. [START_REF] Capet | Upwelling response to coastal wind profiles[END_REF] was poorly simulated, especially between 10°S and 16°S. It is stronger in QSCAT than in ERS, and both satellites have a blind zone near the coast of approximately 25 km and 50 km respectively where no data are available. Note that wind speed was lower in ERS than QSCAT by ~0.5 m s-1 over a large part of the model domain, but that the wind directions were consistent. The differences in wind intensity between ERS and QSCAT are attributed to the use of different operating frequencies, different temporal sampling of the satellites, and gridded products with different spatial resolutions [START_REF] Bentamy | Improvement in air-sea flux estimates derivedfrom satellite observations[END_REF].

Figures 3 d-f show the climatological mean annual cycle of the alongshore wind near the Peru coast. The seasonal cycles were computed over the same time period for ERS and the model (1994)(1995)(1996)(1997)(1998)(1999)(2000), but different time period for QSCAT (2000QSCAT ( -2008)). The observed alongshore winds were strongest during austral winter (July-September) and early spring (October) around 15°S. The model wind climatology was in relatively good agreement with the satellite wind climalogies (Pearson's r>0.85, p<0.01). It reproduced the local maximun values (coastal jets) near 4°S and 15°S which are both seen in QSCAT but not in ERS. ERS did not capture the local maximum at 4°S likely due to its lower spatial resolution. The model overestimated the intensity of the alongshore wind with respect to ERS and QSCAT between 3°S and 17°S, and slightly underestimated it with respect to QSCAT in the equatorial region (0-2°S).

Wind and temperature cross-shore vertical structures

The meridional wind and temperature cross-shore vertical structures in the central Peru (at ~15°S) simulated by the model were compared against the ERA-interim reanalysis data (Figure 4). Reanalysis data showed the coastal jet core located a height of ~250 m and within the first 100 km from the coast, with a wind intensity of ~8 m s⁻ ¹ (Fig. 4a). The coastal jet is capped by a temperature inversion, which base is located at ~600 m above sea level nearshore. The model reproduced relatively well the intensity of the coastal jet core, although with a position much closer to the coast. In overall, the vertical structures of wind and temperature were well reproduced by the model in the first ~1000 m. Note that the steep topography of the Andes is logically represented with greater detail in the WRF model (Fig. 4b).

Surface winds and SST anomalies during El Niño

The model reproduced the spatial structure of the wind anomalies off Peru (north of ~20°S) during November 1997-February 1998 (Fig. 5a) found in the satellite observations (Fig. 1b). Persistent positive (equatorward) alongshore wind anomalies occured from April 1997 to October 1998 (Fig. 5b) in the 5°S-10°S latitude band.

There was a short relaxation period in August 1997 with weak negative anomalies south of 10°S. The strongest positive anomalies (>1.5 m s -1 ) occured between November 1997 and March 1998. Furthermore, strong negative anomalies (<-1 m s -1 ) occured in December 1997 -February 1998 north of 5°S. The general pattern of alongshore wind anomalies from the model was consistent with that from ERS (Fig. 1b). However, some discrepancies were found, such as an overestimation of modelled wind anomalies in May-June 1997 and an underestimation in January 1997. Strong SST positive anomalies (>3°C) were seen between May 1997 and May 1998 along the equator in the Eastern Pacific and along the Peru coast (Fig. 5c), with two main peaks (>4°C) between 5°S and 10°S in July-August 1997and November 1997-April 1998 (Fig. 5d). These peaks were related to the poleward propagation of downwelling coastal-trapped waves, which strongly deepened the thermocline during EN, shutting down the upwelling of cold water [START_REF] Colas | 1997-98 El Niño off Peru: A numerical study[END_REF]. The amplitude of the SST anomalies slightly reduced in September-October 1997 but remained quite strong (>3°C). Note that although the two SST anomaly peaks were of a relatively similar amplitude, only the second peak was synchronous with strong wind anomalies (> 1.5 m s -1 , Fig. 5a) during November 97-April 98, whereas the wind response to the first peak in July-August 1997 was weaker (<1 m s -1 in the model and less than 0.5 m s -1 in ERS; Fig. 1b). Besides, the model also reproduced the negative wind anomalies north of 5°S between November 1997 and April 1998, a period during which SST anomalies were positive.

Alongshore momentum budget

First, in order to investigate the spatial patterns of the forces that induced the wind anomalies off Peru during EN, we computed the anomalies of the meridional component of the forces (involved in the meridional monthly momentum budget, see Eq. 5) in the first layer of the model for the whole domain. Each term was then averaged for the period November 1997-February 1998 (Figs. 6a-d). Second, we computed the monthly alongshore anomalies of the forces (i.e. both zonal and meridional components were used in the projection along the NW-SE direction) and averaged them in a coastal band for the entire EN period (Fig. 6e).

Comparison of anomalies of advection, vertical shear of the meridional turbulent stress (hereafter named vertical mixing term), pressure gradient and Coriolis terms shows that pressure gradient and vertical mixing were the dominant forces (Figs. 6b,c). The pressure gradient anomaly was positive (Fig. 6c), thus accelerated the equatorward wind during EN. Expectedly, vertical mixing anomaly was opposed to the equatorward wind thus negative (Fig. 6b). The two terms almost balanced each other everywhere in the domain, except in the region between 0-5°S where advection anomaly was relatively strong (Fig. 6a) and pressure gradient and vertical mixing anomalies were much weaker. Coriolis term anomalies had much smaller values (Fig. 6d).

Figure 6e shows the time evolution of the monthly anomalies of the surface forces projected in the alongshore direction. The terms were averaged in a one-degree-wide coastal band between 7°S and 15°S (this coastal region was chosen because the coastline is relatively rectilinear and the wind anomalies were relatively homogeneous in space). The pressure gradient and vertical mixing were the main forces during the entire time period. The dates corresponding to pressure gradient maximum (positive anomalies) values coincided with those of vertical mixing minimum (negative anomalies) values (e.g. in June-July and November-December 1997 and April-May 1998).

The advection term was weaker, except in July-September and October-November 1997. The contribution of the Coriolis force was negligible during the entire period. The tendency (ΔV': month to month wind anomaly difference) was almost equal to the sum of all terms, showing that the budget (see Eq. ( 5) in section 2.3) is virtually closed. The small differences in the budget (RMS error of ~6% over the simulation period) come from the acoustic correction (see sections 2.1 and 2.3).

The Consistently with previous studies [START_REF] Muñoz | Dynamics of the low-level jet off the west coast of subtropical South America[END_REF][START_REF] Belmadani | What dynamics drive future wind scenarios for coastal upwelling off Peru and Chile?[END_REF], alongshore pressure gradient anomalies were highly correlated (r=0.8, p<0.01) with alongshore wind anomalies averaged over the coastal band (line red and dashed-black line in Fig. 6e). Although it displays smaller spatial scales than the alongshore wind, the pressure gradient explains relatively well the wind intensification between 5°S and 16°S during the peak of EN but also over limited portions of the coast in June 1997 and August 1998 (Fig. 7a). More specifically, a strong correlation (r=~0.75, p<0.01) between alongshore pressure gradient anomalies and alongshore wind anomalies was found along the coast for latitudes between 4°S and 18°S (Fig. 7b). In this sense, on average over a coastal band (7°S-15°S) ~64% of the temporal variability of the wind anomalies was explained by pressure gradient anomalies, suggesting that vertical mixing behaved like a linear bottom frictional term equilibrating the pressure gradient. This led to the following relation V= In summary, the pressure gradient term played a major role in initiating and terminating the wind anomaly during this EN event. In the next sections we study in detail the processes that drive the pressure gradient increase at the beginning of the event.

Air temperature and humidity contributions to the alongshore pressure gradient

As surface pressure is related to virtual temperature in the air column (see Section 2.4), we examined the relative contributions of temperature and humidity anomalies to the virtual temperature anomalies (VTA) (see Eq. ( 7) in section 2.4) during EN (November 1997-February 1998). The VTA distribution displayed the largest positive anomalies (>4 °K) along the north coast between 4°S and 8°S and in the lowest 300 m (Fig. 8a). Temperature variations dominated VTA (Fig. 8b), while humidity anomalies contributed to at most ~15% of the VTA (north coast around 400 m, Fig. 8c). The humidity anomaly was stronger at 400 m because humidity was higher and nearly constant from the surface up to ~400 m before decreasing progressively with altitude during EN, whereas it progressively decreased with altitude with a relatively constant vertical gradient from the surface to 900m under mean climatological conditions (Fig. 8d). These results show that the stronger temperature increase in the north of Peru was the main driver of the alongshore pressure gradient anomaly during the EN event, while humidity did not play an important role.

This thermally driven pressure gradient was confirmed by the high correlation (r=0.84, p<0.01) between the surface wind anomaly in the coastal band and the alongshore SST gradient (Fig. 9). This correlation was slightly higher (0.87) with a 1 month lag (when the SST gradient leads the wind). This suggests that the anomalously warm surface ocean forced the low atmosphere by heating the air column more in the north than in the south, thus generating the pressure gradient that drove the equatorward wind anomaly.

Downward mixing of momentum during EN

Due to the air warming and humidification associated with the presence of anomalously warm surface waters in the nearshore region, shallow convection was enhanced. In the coastal band, the planetary boundary layer height (PBLH) increased by ~200 m (~50%) around June 97 and by ~100-150m (~100%) between December 1997-March 1998 (Fig. 10). The PBLH increases were in phase with the SST anomalies peaks, and slightly stronger in the north than in the south, in agreement with the SST spatial changes (not shown).

Given this PBL variability, a potential mechanism for the wind intensification could be associated with the increase of turbulence in the PBL during EN, which may generate a more efficient downward vertical flux of momentum [START_REF] Wallace | The influence of sea surface temperature on surface wind in the Eastern[END_REF]). However, turbulent mixing tended to decelerate the wind even more during the warm EN phases (Fig. 6e). To further investigate why [START_REF] Wallace | The influence of sea surface temperature on surface wind in the Eastern[END_REF]'s mechanism can not explain the EN wind increase, we examined the vertical profiles of alongshore wind, turbulent vertical mixing coefficient, momentum budget forces contribution, and temperature during climatological and EN conditions in November 1997-February 1998 (Fig. 11). The wind intensification during EN occurred between the sea surface and ~1600 m (black line in Fig. 11a). The wind maximum (~8 m s -1 ) shifted from ~300 m in mean climatological conditions to ~500 m. There was a decrease of wind shear (dV/dz) below ~500m during EN, mainly due to the velocity increase at the surface (~20%) and a virtually unchanged velocity at 300 m (black line in Fig. 11a). Turbulent vertical mixing coefficient (Kz) increased almost twofold during EN reaching a maximum of ~40 m².s -1 at 200 m (~15 m 2 s -1 in climatological conditions, blue line in Fig. 11a). The impact on the turbulent stress (τ=Kz.dV/dz) of the wind shear decrease and turbulent vertical mixing (Kz) increase was such that the momentum vertical mixing term (dτ/dz) reduced ( it is negative in mean climatological conditions, blue line in Fig 11b and12b)

during EN conditions (Fig. 6e; blue line in Fig. 11c). This also shows that the EN wind intensification near the surface was not driven by downward mixing of momentum. The pressure gradient (positive in mean climatological conditions, red line in Fig. 11b and12b) increase was maximum at the surface and decreased with height during EN conditions (red line in Fig. 11c). The Coriolis and advection terms did not change much during EN (magenta and green lines in Fig. 11c), but they are relatively important for the budget in mean climatological conditions (magenta and green lines in Fig. 11b and12b). Note that the pressure gradient was strong at surface and in upper layers during EN. Air temperature decreased between the surface and 1600 m, showing no temperature inversion in this period (austral summer) in mean climatological or EN conditions (Fig. 11d).

This time period is contrasted with the period July-August 1997, during which the surface wind anomalies were weak (Figs. 5a and12a) in spite of SST anomalies of the same order (~4°C) as during November 1997-April 1998 (Fig. 5b). Between the surface and 800 m the wind speed decreased with respect to mean climatological conditions, and did not change in the upper layers (800-1600m). The wind shear decreased between the surface and 400m (black line in Fig. 12a), mainly due to the velocity decrease between 300-400 m (~10%) and a virtually unchanged velocity at surface. Mixing coefficient (Kz) increased by almost 70% in this period, reaching a maximum of ~74 m².s -1 at 250 m (~45 m 2 s -1 in climatological conditions at 220 m, blue line in Fig. 12a). The changes in wind shear and turbulent vertical mixing (Kz) in this period counteracted such that the momentum vertical mixing term did not change significantly (blue line in fig 12c), thus compensation . The wind decrease below ~600 m was likely driven by advection, which decreased below 500 m (green line in Fig. 12c) whereas the pressure gradient change remained positive (red line in Fig. 12c). On the other hand, the pressure gradient vertical shear between the surface and ~600 m (Fig. 12c) was stronger than in November 1997-February 1998 period (Fig. 11c). Note also the well-marked temperature inversion in July-August 1997 (Fig. 12d). The effect of this particular vertical structure on the pressure gradient will be discussed below (Sect. 4.2).

Impacts of large scale atmospheric forcing and local SST forcing

In There was a slightly negative anomaly in September 1997, which coincided with the slight decrease of SST anomaly in September-October 1997 (Fig. 5b). Note that the wind anomaly was negligible in CRTL in Jul-Aug 1997, in spite of a strong SST anomaly (~3-4 °C, Fig. 5b). This can be explained by the large scale signal, which forced a decrease of the equatorward wind (negative anomalies of ~ -1 m s -1 in July-September 1997 in BRY-EN) which compensated the wind increase (~ 0.5-1 m s -1 in SST-EN) forced by the anomalously warm SST.

In the first warm period (July-August 1997), the intensification of the wind driven by the SST forcing (SST-EN experiment) was compensated by the large scale signal (BRY-EN) not only at the surface but also as high as 1400 m (Fig. 13b). During the second period (November 1997-February 1998), boundary conditions (BRY-EN simulation) did not have a strong effect. The wind was modified very little in the vertical with respect to climatological conditions. The wind intensification below 200 m was fully forced by the anomalous SST (Fig. 13c).

Thus, the BRY-EN experiment showed that the large-scale atmospheric signal propagating into the Peru region through the open boundaries could mitigate (or enhance) the coastal wind anomalies during EN. The SPA center was located around its mean climatological position into both periods (red lines in Figs 14 a,b). However, during

March-September 1997, the SPA was weaker (-1-2 hPa anomaly) than in mean climatological conditions (Fig. 14a). This weakening should produce negative anomalies for the surface atmospheric pressure off Peru, which likely contributed to a decrease of the alongshore wind off Peru in the BRY-EN simulation. Indeed, the maximum surface pressure in the SPA was well correlated (0.67) with BRY-EN wind anomalies during the EN period. In contrast with the March-September 1997 period, the SPA was slightly more intense in October 1997-March 1998 (Fig. 14b) and September 98 (not shown), favoring an intensification of the coastal wind during these time periods due to SST forcing (Fig. 13a).

Discussion and conclusions

Summary

A regional atmospheric model was used to investigate the physical processes driving the wind intensification off the Peru coast during the 1997-1998 EN event. As anomalously warm waters accumulated near the coast, the equatorward coastal wind increased by ~1-1.5 m s -1 during 5-6 months (up to ~40 % increase with respect to the climatological mean over the 1994-2000 period). Simulated surface wind anomalies during EN were in good agreement with observed wind anomalies. A momentum balance analysis showed that the coastal wind intensification was mainly driven by the enhancement of the alongshore pressure gradient. Vertical mixing tended to counterbalance the alongshore pressure gradient, leading to a quasi-equilibrium between the alongshore pressure gradient and the frictional force, consistently with previous modeling studies in the region [START_REF] Muñoz | Dynamics of the low-level jet off the west coast of subtropical South America[END_REF]Belmadani et al., 2013). The enhancement of the alongshore pressure gradient occured because the atmospheric pressure decreased more north (~6°S) than south (~14°S), in association with the larger increase of SST, air temperature and humidity off northern Peru. Surface warming induced an increase of the height of the PBL of up to two and half times and of the vertical turbulent mixing coefficient (Kz=τ/(dV/dz)) of up to three times their values in mean climatological conditions. However vertical mixing of momentum (dτ/dz) remained negative and was stronger (in absolute value) during EN that in mean climatological conditions, thus did not accelerate the equatorward wind.

The back-pressure effect during EN

The alongshore pressure gradient change was strong at surface and its vertical structure varied during the EN period: the pressure gradient change was strong between the surface and ~1000m when there was no temperature inversion (e.g. in November 1997-February 1998;Figs. 11b,c). In contrast, it became negligible above 600m in the presence of a marked temperature inversion (e.g. in July-August 1997; Figs. 12b,c). [START_REF] Hashizume | Direct observations of atmospheric boundary layer response to SST variations associated with tropical instability waves over the Eastern Equatorial Pacific[END_REF] showed that the so-called "back-pressure effect", a mechanism compensating the surface pressure gradients, was strong in the case of a marked temperature inversion above the PBL. [START_REF] Oerder | Mesoscale SSTwind stress coupling in the Peru-Chile current system: Which mechanisms drive its seasonal variability?[END_REF] found that this effect reduced significantly the intensity of the surface pressure gradient above mesoscale SST positive anomalies in the Peru region. This is also likely the case in the Peru coastal region in our simulation during the period July-August 1997 (Fig. 12c). Note that temperature inversion is related to the strong atmospheric subsidence in the region led by the SPA (e.g. [START_REF] Haraguchi | Inversions over the tropical eastern Pacific ocean[END_REF]).

In conclusion, it is likely that two conditions, a SST anomaly alongshore gradient and a weak (or absent) temperature vertical inversion would be necessary to drive a strong surface wind anomaly in the coastal region.

The SST gradient drives the pressure gradient, and the weak temperature inversion allows shallow convection to develop without triggering any "back-pressure effect". [START_REF] Bakun | Greenhouse gas, upwelling favourable winds, and the future of upwelling systems[END_REF] suggested that the increase of humidity over coastal land during EN could enhance the local greenhouse heating effect, thus increasing land temperature more than SST. This thermal contrast would lead to the presence of lighter air over land than over sea, an intensification of the cross-shore pressure gradient and associated alongshore geostrophic wind. Similarly, Enfield (1981) suggested that a land-sea thermal contrast may occur due to an increase of downward solar radiation over land, which would be due to a reduction of cloudiness during EN. As [START_REF] Bakun | Greenhouse gas, upwelling favourable winds, and the future of upwelling systems[END_REF][START_REF] Co;2. Enfield | Thermally-driven wind variability in the planetary boundary layer above Lima, Peru[END_REF] suggested, humidity and incoming solar radiation increased over land during EN in our simulations (green and red lines respectively; Fig. 15). However, this increase in humidity and short wave radiation reaching the land surface did not support a strengthened land-sea thermal contrast. Indeed, the simulated land-sea contrast, which was positive (i.e. air temperature was higher over land than over sea) in mean climatological conditions (not shown), decreased by ~100% during EN (blue line in Fig. 15) as the air temperature over sea increased much more than the air temperature over land (not shown). Note that a clear relation between land-sea thermal contrast and alongshore winds off Peru has not been demonstrated in previous studies. Using a set of GCM simulations with a spatial resolution of ~50 km in the Peru-Chile region, Belmadani et al. (2013) found that the land-sea thermal contrast (dT/dx) increased off Peru in scenarios of climate change: at 8°S, dT/dx was ~4.5 10 -2 K km -1 in preindustrial climate conditions, and reached ~6.5 10 -2 K km -1 in 4xCO2 climate conditions, thus increased by ~40% (see Figure 8c in Belmadani et al. 2013).

Land-sea thermal contrast

In spite of this strong increase, the alongshore wind decreased moderately (~10%) off Peru. This reduction was driven by a decrease of the alongshore pressure gradient associated with a poleward displacement of the SPA in a warmer climate (Belmadani et al. 2013 and references herein). This and our results suggest that an increase of the land-sea thermal gradient may not play a strong dynamical role for the alongshore wind, at least for the range of horizontal resolution (~25 km in the present study and ~50 km in Belmadani et al. 2013) explored in our model simulations. Note that these resolutions do not allow to represent the coastal terrain located between the Andes and the ocean with more than a few grid points.

In addition, our model simulated poorly the mean downward shortwave radiation associated with the cloud cover, a well-known problem in models of the Southeast Pacific lower troposphere (e.g. Wyant et al. 2010).

Despite this bias, the model reproduced reasonably well the surface air temperature distribution and, more importantly for the purpose of the present study, the air temperature anomalies at surface during EN, in agreement with reanalysis data (not shown). The modelled air temperature increased more over sea than over the land during EN associated with the strong SST warming, thus reducing the land-sea thermal contrast. However, the land-sea thermal contrast may impact the wind at scales smaller than those resolved by the present model (e.g. [START_REF] Co;2. Enfield | Thermally-driven wind variability in the planetary boundary layer above Lima, Peru[END_REF]. This process remains to be investigated using higher resolution model experiments in future work.

Dynamical processes during other EN events

Due to the model computational cost, our simulations were performed for a relatively short time period (1994)(1995)(1996)(1997)(1998)(1999)(2000) including only one EN event. In order to evaluate if the same dynamical processes were active during other events, we performed similar diagnostics using the ERA-Interim reanalysis data over the period 1979-2016. This data has a lower spatial resolution (~80 km) than our regional model, but it can give hints of the processes at stake during EN events. The wind anomaly (in a coastal band of ~160 km and between 7°S and 15°S) for the 1997-98 EN was ~0.6-0.8 m s -1 , less than in our model ( ~1.0-1.2 m s -1 for an average over 6 WRF grid points). We found relatively strong wind anomalies during most EN events, in particular in 1982-83 (~0.8 m s -1 ), 1987-88 (~1 m s -1 ), 1992-93 (0.7 m s⁻ ¹), 2015-16 (0.8 m s⁻ ¹) (Fig. 16). In agreement with our analysis, these four events were associated with positive alongshore SST gradient anomalies (e.g. ~0.15 10⁻ ² °C km⁻ ¹ in 1982-83). There were also EN events with relatively weak wind anomalies (i.e. 2002-2003). Note that strong SST gradient anomalies occurred during relatively short time periods in 1993, 2002 and 2008, and were not associated with positive wind anomalies. We may conjecture that other processes such as a compensation by the large scale forcing through a modification of the SPA or the presence of temperature inversion with a backpressure effect may be active during these periods.

Local air-sea coupled processes during EN

Local air-sea coupled interactions not investigated in the present study may also play a role during EN. First, the increase of humidity in the north of Peru leads to intense precipitation on land and over the nearshore ocean [START_REF] Takahashi | The atmospheric circulation associated with extreme rainfall events in Piura, Peru, during the 1997-1998 and 2002 El Niño events[END_REF], which may enhance the ocean surface stratification. This may mitigate wind-driven oceanic vertical mixing in the north and thus help maintaining the anomalous alongshore SST gradient driving the coastal wind. On the other hand, the stronger coastal wind (Figs 1b and 5a), SST (Fig. 5b) and humidity (Fig. 8c) anomalies in the north would increase evaporation and thus cool the ocean more efficiently than further south.

This effect may mitigate the SST gradient and thus reduce the wind anomaly. Studying such feedbacks, which

were not taken into account in our forced atmospheric model framework, is beyond the scope of the present Table 1. Parameterizations used in WRF model for the simulations.

Processes Scheme Reference

Shortwave Radiation Dudhia scheme [START_REF] Dudhia | Numerical study of convection observed during the winter monsoon experiment using a mesoscale two-dimensional model[END_REF] Longwave Radiation RRTM scheme [START_REF] Mlawer | Radiative transfer for inhomogeneous atmosphere: RRTM, a validated correlated-k model for the long-wave[END_REF] Microphysics WRF Single-Moment 6-class scheme [START_REF] Hong | The WRF single-moment 6-class microphysics scheme (WSM6)[END_REF] Cumulus Betts-Miller-Janjic scheme [START_REF] Janjic | The step-mountain eta coordinate model: further developments of the convection, viscous sublayer, and turbulence closure schemes[END_REF] Surface Layer MYNN surface layer [START_REF] Nakanishi | Development of an improved turbulence closure model for the atmospheric boundary layer[END_REF] Land Surface Noah Land Surface Model [START_REF] Chen | Coupling an advanced land surface hydrology model with the penn state-NCAR MM5 modeling system. Part ii: Preliminary model validation[END_REF] Planetary Boundary Layer MYNN Level 2.5 PBL [START_REF] Nakanishi | Development of an improved turbulence closure model for the atmospheric boundary layer[END_REF] 8. [1994][1995][1996][1997][1998][1999][2000].

Figure 15. Anomalies of air humidity at 2 meters (in g kg-¹, green line), downward shortwave radiation at sea surface (in W m -2 , red line) and land-sea air temperature gradient at 2 meters (in 10 -2 °C km, blue line).

Anomalies of air humidity and shortwave radiation were taken at the model land grid points closest to sea.

Temperature gradient was computed as the difference between the model land grid point closest to sea and the sea grid point closest to land at each latitude. All variables were averaged between 7°S and 15°S.

-1 -1 -1 - 1 -1 - 1 -1 -1 -1 -1 -1 -1 -1 -1 -1 -1 -1 -1 -1 -1 -1 -1 -1 -1 -1 -1 20 o S

  wind intensification during EN occurred during different phases (Fig. 6e, dashed-black line). It began in March 1997 (ΔV'~1 m s⁻ ¹ from March to April 1997) and was maintained until June 1997, due to a positive equatorward pressure force stronger than the sum of the other (negative) terms. The wind anomaly then strongly decreased (ΔV'~-1.2 m s -1 from June to July 1997) due to a negative advection of momentum between June and August 1997, which was related to the large scale forcing. A similar momentum budget for the BRY-EN experiment confirmed this large scale modulation (figure not shown). The wind acceleration became strong again in September 1997 (ΔV'~1.5 m s -1 from September to October), when advection weakened and the pressure gradient dominated the other forces. The tendency was weak from January to May 1998. In the later phase of EN, the wind anomaly was strongly reduced from May to July 1998, due to a decrease of the pressure gradient.

  linear friction coefficient and ρ the surface air density. From our model results, c was ~ 4 10 -5 s -1 , thus comparable to[START_REF] Muñoz | Dynamics of the low-level jet off the west coast of subtropical South America[END_REF]'s estimation for the Chile central coast (c ~ 5 10 -5 s -1 ). The steep orography of the Andes precluded the development of an anomalous cross shore flow, thus the alongshore pressure gradient can not be equilibrated by the Coriolis force. Consequently the alongshore flow built up until the frictional force (i.e. vertical mixing term) balanced the alongshore pressure gradient.

  this subsection, we analyzed the model sensitivity experiments to study the respective roles of the large scale atmospheric signal (BRY EN experiment, forced by 1997-1998 EN boundary conditions and neutral SST forcing from the years 1994-1995) and of the SST local forcing (SST-EN experiment, forced by EN SST forcing and 1994-1995 neutral boundary conditions, see section 2.5) in driving the EN wind anomalies.

Figure

  Figure 13a displays the alongshore surface wind anomalies (averaged in the coastal band between 7°S and 15°S) for the CRTL, BRY-EN, SST-EN experiments. The large scale signal (BRY-EN simulation) during EN induced strong negative wind anomalies (~ -1 m s -1 ) between March and September 1997, and moderate (<1 m s -1 ) positive anomalies between October 1997 and March 1998. In contrast, the model forced by the 1997-98 SST forcing (SST-EN) simulated persistent positive wind anomalies between May 97 and August 97 (reaching ~1 m s -1 in August 97) and between November 1997 and May 1998 (peaking at ~1.5 m s -1 in January-February 98).

  Equatorial Pacific: seasonal and interannual variability. JClim 2:1492Clim 2: -1499Clim 2: . doi:10.1175Clim 2: /1520Clim 2: - 0442(1989))002<1492:TIOSST>2.0.CO;2 Wyant MC et al (2010) The PreVOCA experiment: modeling the lower troposphere in the Southeast Pacific. Atmos Chem Phys 10:4757-4774. doi:10.5194/acp-10-4757-2010 Wyrtki K (1975). El Niño: the dynamic response of the equatorial Pacific to atmospheric forcing. Journal of Physical Oceanography 5:572-584 Xue H, Wang Y, Xie SP (2004) Effects of the andes on eastern pacific climate: a regional atmospheric model study. J Clim 17:587-602 Zeng X, Beljaars A (2005) A prognostic scheme of sea surface skin temperature for modeling and data assimilation. Geophysical Research Letters 31:L14605. doi: 10.1029/2005GL023030 Zuta S, Guillén O (1970) Oceanografía de las aguas costeras del Perú. Boletin Instituto del Mar del Perú 2

Figure 1 .

 1 Figure 1. a) ERS wind anomalies (in m s -1 ) off Peru and Northern Chile during El Niño conditions in November-February 1997/1998. Arrows mark the direction of the monthly wind anomalies. b) Time-latitude diagram of ERS alongshore monthly wind anomalies off the Peru coast. The wind average was computed within a 100 kmwide coastal band and a 3-month running mean was applied. Positive values indicate equatorward wind anomalies.

Figure 2 .

 2 Figure 2. South East Pacific model domain and Peru nested domain used in the WRF simulations. Color shading indicates model topography (in meters) above sea level for the small domain.

Figure 3 .

 3 Figure 3. Mean surface wind (in m s -1 ) in 2000 from a) ERS satellite, b) WRF model and c) QSCAT satellite. Mean annual cycle of the alongshore wind (averaged in a 100-km-wide coastal domain) from d) ERS and e) WRF over the period 1994-2000, and from f) QSCAT over the period 2000-2008.

Figure 4 .

 4 Figure 4. Vertical structure of the mean meridional wind (shading, in m s -1 ) and air temperature (black contours, in °C) at 15°S from a) ERA-Interim and b) WRF model. Data was averaged for the 1994-2000 period.

Figure 5

 5 Figure 5. a) Mean surface wind anomalies (in m s -1 ) from WRF (over November 1997-February 1998) and b) time-latitude diagram of WRF alongshore wind anomalies. c) Mean sea surface temperature (SST) anomalies (in °C) from OISST (over the same time period in a)) and d) time-latitude diagram of alongshore SST anomalies.

Figure 6 .

 6 Figure 6. Anomalies of the meridional component of the surface forces contribution (in m s -1 ) during El Niño 1997/1998: a) advection (V-ADV'), b) vertical turbulent mixing (V-MIX'), c) pressure gradient (V-PGF'), and d) Coriolis force (V-COR'). Anomalies were computed with respect to a climatology over 1994-2000 and averaged over November 1997-February 1998. e) Time series of the monthly anomalies of the alongshore component of the surface forces contribution in 1997-1998 (in m.s⁻ ¹, with scale defined by the left-hand side y axis). The forces were averaged in a 100-km coastal band between 7°S and 15°S. Gray, black and black-dashed lines

Figure 7

 7 Figure 7. a) Alongshore anomalies of pressure gradient (in m s -1 , shading) and coastal wind (in m s -1 , black contours) between the equator and 20°S. b) Latitudinal variation of the correlation between alongshore pressure gradient anomalies and alongshore coastal wind anomalies.

Figure 8 .

 8 Figure 8. Alongshore vertical sections of the mean monthly anomalies (during November 1997-February 1998) of a) virtual temperature (in °C) and b) temperature and c) humidity contributions (in °C) to the virtual temperature anomaly. Note the different scales for temperature and humidity contribution. d) Vertical profiles of humidity in the northern coastal region (4°S-8°S) during November 1997-January 1998 (full line) and mean climatological conditions (dashed line).

Figure 9 :

 9 Figure9: Time evolution of alongshore SST gradient (in °C/25km, positive equatorward, red line) and wind (in m s -1 , black line) anomalies. Anomalies were smoothed using a 3-month running mean, and averaged between 7°S and 15°S and within 100 km from the coast.

Figure 10 .

 10 Figure 10. Mean Planetary Boundary Layer Height (in meters, PBLH) off Peru. PBLH was averaged between 7°S and 15°S and within 100 km from the coast. Red line marks PBLH during 1997-1998, and black dashed line the climatology. Error bars indicate standard deviation from the mean climatological values.

Figure 11 .

 11 Figure 11. Vertical profiles of a) alongshore wind (in m s -1 ), b) alongshore forces in mean climatological conditions (in m s⁻ ¹), c) anomalies of alongshore forces in Niño conditions (in m s -1 ) and d) air temperature (in °C). Averages were computed over November 1997-February 1998, between 7°S and 15°S, and within 100 km from the coast. Full and dashed lines in a) and c) correspond to El Niño and mean climatological conditions, respectively.

Figure 13 :

 13 Figure 13: a) Time series of coastal alongshore wind anomalies (in m s -1 ) from CRTL (black line), SST-EN (red line), BRY-EN (blue line) experiments. Mean alongshore wind profiles (in m s -1 ) in b) July-August 1997 and c) November 1997-February 1998. Black-dashed, black, red and blue lines mark the climatological (CLIM) ,CRTL, SST-EN and BRY-EN profiles, respectively.

Figure 14 .

 14 Figure 14. Mean (red contours) and anomalous (shading) sea level pressure (in hPa) during a) March-September 1997 and b) October 1997-March 1998. White contours mark pressure climatological values(1994)(1995)(1996)(1997)(1998)(1999)(2000).

Figure 16 .

 16 Figure 16. Time evolution of alongshore SST gradient (in °C/80 km, positive equatorward, red line) and alongshore wind (in m s -1 , black line) anomalies for ERA-Interim reanalysis over 1979-2016. Anomalies were averaged in a 160-km-wide coastal band and between 7°S and 15°S. Grey vertical bands indicate El Niño periods.

Figure 5 Figure 9

 59 Figure 5

Table 1 .

 1 Parameterizations used in WRF model for the simulations.

	Processes	Scheme	Reference
	Shortwave Radiation	Dudhia scheme	Dudhia (1989)
	Longwave Radiation	RRTM scheme	Mlawer et al. (1997)
	Microphysics	WRF Single-Moment 6-class	Hong and Lim (2006)
		scheme
	Cumulus	Betts-Miller-Janjic scheme	Janjic (1994)
	Surface Layer	MYNN surface layer	Nakanishi and Niino (2009)
	Land Surface	Noah Land Surface Model	Chen and Dudhia (2001)
	Planetary Boundary Layer	MYNN Level 2.5 PBL	Nakanishi and Niino (2009)

Table 1

 1

study. These questions, which can be addressed using a regional high resolution, ocean-atmosphere coupled model (e.g. Oerder at al. 2016), will be the purpose of future studies.

Acknowledgments

This research is part of A. Chamorro's PhD thesis, funded by the fellowship from CIENCIACTIVA/CONCYTEC-PERU at the University Pierre and Marie CURIE of France. It is also part of the IDB project PE-G1001(Adaptation to Climate Change of the Fishery Sector and Marine-Coastal Ecosystem), and it is a contribution to the cooperative agreement between the Instituto del Mar del Peru (IMARPE) and the Institut de Recherche pour le Developpement (IRD) and the LMI DISCOH. The simulations were performed on the supercomputer Curie from the GENCI at the CEA (projects 2011040542, 2012061047 and 2014102286).

Francois Pinsard is acknowledged for her help in the making of lateral boundary forcing for the regional atmospheric model. Francis Codron and Clémentine Junquas are acknowledged for useful discussions.