Mar 1, 1997 - Laboratoire d'Ocйanographie Dynamique et de Climatologie ... Paris, France. Received 9 ... 1995; McPhaden, 1999; Boulanger and Menkes, 1999; ...... (CNRS/UPMC/IRD), Universitй Pierre et Marie Curie, Tour 26, 4eme.
JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 107, NO. C12, 8015, doi:10.1029/2001JC000841, 2002
Ocean response to the March 1997 Westerly Wind Event Matthieu Lengaigne, Jean-Philippe Boulanger, Christophe Menkes, Sebastien Masson, Gurvan Madec, and Pascale Delecluse Laboratoire d’Oce´anographie Dynamique et de Climatologie (CNRS/UPMC/IRD), Universite´ Pierre et Marie Curie, Paris, France Received 9 February 2001; revised 9 October 2001; accepted 16 October 2001; published 26 November 2002.
[1] An Ocean General Circulation Model is used to investigate the oceanic response to
the March 1997 Westerly Wind Event that is suggested to have played an important role in the onset of the 1997–1998 El Nin˜o. Our results point out three distinct impacts. First a strong wind-forced downwelling Kelvin wave propagates eastward generating sea surface temperature anomalies up to 1C and large subsurface temperature and zonal current anomalies, mainly located in the core of the thermocline. Second the northward and westward extension of this wind event is responsible for a surface advection of cold waters from 130E–5N to the equator. Third it generates large zonal surface currents at the eastern edge of the warm and fresh pool by a nonlinear interaction between the windforced surface jet and the local thermohaline front. Salinity through both its contribution to the local zonal pressure gradient at the front and the barrier layer effect is crucial in the occurrence of this nonlinear mechanism. The fast displacement of the front (2000 km in a month) together with the cooling in the western Pacific is likely to be responsible for the eastward displacement of atmospheric deep convection and eastward winds observed in April–June 1997 and thus to have played a major role in initiating the El Nin˜o of the INDEX TERMS: 4522 Oceanography: Physical: El Nin˜o; 4572 Oceanography: Physical: Upper century. ocean processes; 4504 Oceanography: Physical: Air/sea interactions (0312); 4512 Oceanography: Physical: Currents; KEYWORDS: Westerly Wind Events, El Nin˜o, equatorial oceanic waves, nonlinear physics, Pacific warm pool Citation: Lengaigne, M., J.-P. Boulanger, C. Menkes, S. Masson, G. Madec, and P. Delecluse, Ocean response to the March 1997 Westerly Wind Event, J. Geophys. Res., 107(C12), 8015, doi:10.1029/2001JC000841, 2002.
1. Introduction [2] Whereas easterly trade winds prevail over most of the equatorial Pacific, weak reversing monsoon winds seasonally replace the trade winds over the western Pacific warmpool during boreal winter. Superimposed to the seasonal signal, energetic synoptic scale westerly wind events (WWEs) are often observed over the warm-pool [Giese and Harrison, 1991; Harrison and Vecchi, 1997]. They are associated with Madden Julian Oscillations and tropical cyclone formation [Hendon et al., 1999; Harrison and Giese, 1991]. In the ocean, these WWEs force equatorial downwelling Kelvin waves that cross the Pacific, induce an eastward displacement of the warm-pool, and may contribute to the onset of El Nin˜o events [Picaut and Delcroix, 1995; McPhaden, 1999; Boulanger and Menkes, 1999; Slingo and Delecluse, 1999]. The ocean response to WWEs has been analyzed in many studies using observations [Delcroix et al., 1993; Feng et al., 1998; Johnson et al., 2000], or models [Giese and Harrison, 1993; Kindle and Phoebus, 1995; Maes et al., 1998; Richardson et al., 1999; Kessler and Kleeman, 2000]. They all show a strong sea surface temperature response: a local cooling in the west Copyright 2002 by the American Geophysical Union. 0148-0227/02/2001JC000841
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and a remote heating in the east. None of them have focused on the role of salinity on that response, although it may be important [Roemmich et al., 1993]. Indeed, on long term average, the western Pacific is characterized by weakly positive heat fluxes [Oberhuber, 1988] and heavy precipitation. These atmospheric conditions often result in the presence of a barrier layer (BL) over the warm-pool [Lukas and Lindstrom, 1991], an area where the mixed layer depth is controlled by the haline stratification rather than the thermal structure. The BL affects both the SST and surface current responses to WWEs [Vialard and Delecluse, 1998b]. By restraining exchanges with cooler waters from below, the BL retains heat and momentum in the upper layer of the warm-pool and increases the magnitude of surface jets. Moreover, Roemmich et al. [1993] suggests that the existence of a zonal salinity gradient may also significantly contribute to drive eastward jets. Recently Boulanger et al. [2001] emphasizes the role of a nonlinear interaction between the surface jet and the thermohaline front at the eastern edge of the warm-pool that strongly increases the eastward displacement of the edge of the warm-pool. [3] In this study, we focus on the onset of the 1997 – 1998 El Nin˜o which has been extensively observed through both satellite and in situ measurements. It is associated with a series of intense WWEs occurring between December 1996 and June 1997. The strongest in terms of intensity (up to 2
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dyn/cm2), fetch (more than 30 longitude) and duration (about a month) was observed from mid-February to midMarch. It has been suggested that this event might have had a strong impact on the 1997 – 1998 El Nin˜o onset in displacing the warm-pool eastward and thus in potentially initiating ocean-atmosphere feedbacks leading to the El Nin˜o of the century [McPhaden, 1999; McPhaden and Yu, 1999; Boulanger and Menkes, 1999]. Our aim is to investigate the basin scale ocean response to the March WWE using an Ocean General Circulation Model. We will pay particular attention to the role of salinity on the ocean response during this event. [4] This paper is organized as follows: section 2 presents the modeling framework and a brief model validation. Section 3 describes the ocean response to the March 1997 WWE, first in the central and eastern Pacific where a strong wind-forced Kelvin wave propagates, second in the western Pacific warm-pool where the WWE occurs, and third at the eastern edge of the warm and fresh pool. Section 4 discusses the specific impact of salinity on the ocean response to the March WWE. A conclusion is given in section 5.
2. Reference Experiment [5] The primitive equation model employed in this study is the OPA8 OGCM [Madec et al., 1998] with a free surface formulation [Roullet and Madec, 2000] in its global configuration ORCA. The horizontal resolution is 2 in longitude and varies in latitude from 0.5 at the equator to 2cos(f) poleward from the tropics. In order to remove the North Pole singularity, two poles are introduced in the Northern Hemisphere grid (one over the Canada and the other over Siberia). The departure from a geographic mesh starts at 20N, and the stretched part of the mesh is constructed following the semi-analytical method of Madec and Imbard [1996]. There are 31 vertical levels with the highest resolution (10 m) in the upper 150 meters and the time step is 1h36’. The Jackett and McDougall [1995] equation of state is used. It provides density as a function of potential temperature, salinity and pressure. Vertical mixing coefficients are computed from a 1.5 turbulent closure scheme [Blanke and Delecluse, 1993] which permits an explicit formulation of the mixed layer as well as a minimum diffusion in the thermocline. An isopycnal mixing is used on both momentum and active tracers with an eddy coefficient of 2000 m2 s1 (M. Lengaigne et al., Sensitivity of the tropical Pacific Ocean to isopycnal mixing, manuscript submitted to Journal of Physical Oceanography, 2001). [6] The 1993 –1999 surface forcing fields are given by the ERS1-2 weekly wind stress [Bentamy et al., 1996] and the daily NCEP reanalysis heat and freshwater fluxes [Kalnay et al., 1996]. The strong coupling between surface ocean heat flux and SST is approximated by a local negative feedback term toward the observed interannual SST: Qref ¼ Qncep þ
dQ SSTref SSTobs dT
ð1Þ
where Qref is the net heat flux at the air-sea interface; SSTref is the model SST, SSTobs is the observed SST of Reynolds and Smith [1994] observed SST, Qncep is the NCEP heat flux, and
dQ/ dT is a restoring coefficient taken equal to 40 W m2 K1, a typical value of the tropical oceans. There is no feedback term toward the observed sea surface salinity (SSS). [7] The reference experiment (REF) starts after a five year spin up phase obtained with a climatological forcing derived from the above forcing fields. It covers the 1993 – 1999 period. An exhaustive validation of the Pacific Ocean’s response is beyond the scope of the paper. The model behavior and its comparison to observations (altimetry, TAO moorings) are quite similar to those of Vialard et al. [2001] and Radenac et al. [2001] who used the same model in a Pacific configuration forced over the same period by the same wind product but thermohaline fluxes from ERA15 reanalysis [Gibson et al., 1997]. The change in forcing fields does not significantly affect the model-data comparison. The model succeeds in reproducing the basin wide structures of currents, sea level, and temperature, and accurately simulated Kelvin wave activity (which is of particular importance in our study). As an example, Figure 1 shows a comparison to TOPEX/POSEIDON sea level anomalies. As in Vialard et al. [2001] the correlation is quite high (up to 0.9 in the 5N– 5S band) and the root mean square (RMS) differences are far below the RMS for each dataset (around 4 cm in the 5N–5S band).
3. Oceanic Response to the March 1997 WWE 3.1. March 1997 WWE [8] The March event is the second of a series of strong WWEs observed during the December 1996– June 1997 period. The first WWE occurs near 150E at the end of December 1996 but its effect on the central and eastern Pacific is counteracted by the presence of strong easterlies during the first two months of 1997. These easterlies confine the warm and fresh pool to the west of the dateline [Boulanger and Menkes, 2001; Vialard et al., 2001; Radenac et al., 2001]. Contrary to the majority of WWEs [Harrison and Vecchi, 1997], the March 1997 event lasted more than one month, from mid-February to mid-March 1997 (Figure 2a). Its maximum amplitude extends zonally over about 30 from 135 to 165E, and meridionally over more than 20 from a few degrees north of the equator to about 15S (Figure 2). At the equator, this particularly intense WWE reaches its maximum around the 10th of March (Figure 2c). At that time, the zonal wind is maximum to the south of the equator (more than 2 dyn/cm2) and the meridional component is essentially negative with a maximum amplitude around 0.1 dyn cm2. This event can be considered as an S-type event in the classification of Harrison and Giese [1991]. [9] In order to investigate the oceanic response of the March 1997 WWE, we performed a sensitivity experiment referred to as NWE (No Wind Event). In that experiment, the March 1997 WWE was eliminated from the ERS wind stress data by simply zeroing all positive zonal wind stress signal over the WWE duration in the region encompassed by the WWE (i.e. 10N–10S, 120E –180E). The heat and freshwater fluxes in NWE are exactly the same as in REF. To that purpose, the net surface heat flux of REF experiment has been stored and is applied to NWE without any feedback term. Therefore the difference REF-NWE is only related to a change in the wind stress field.
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Figure 1. Comparison of TOPEX/POSEIDON and REF experiment sea surface anomaly for the 1993 – 1999 period. (a) correlation and (b) Root Mean Square difference. Contour interval is 0.1 and 1 cm, respectively.
[10] Since the atmospheric forcing in NEW remains the same as in REF after the March WWE, the NWE ocean state will tend toward that of REF within a few months. Indeed, these are almost identical after 8 months. Thus one should not study the NWE experiment over a too long period because, in reality, the atmospheric response is certainly different with and without the March WWE whereas we keep it constant by means of methodology employed. Therefore in the following, the effect of the March 1997 WWE is only studied until May 1997. Figure 3 shows equatorial Hovmueller plots of SST, SSS and zonal surface currents for REF and their differences with NWE from February to May 1997. The oceanic response to the
March 1997 WWE can be split into three main areas : (1) the central and eastern Pacific where a wind-forced Kelvin wave propagates; (2) the far western Pacific where a local response to the WWE occurs ; and (3) the dateline region where a specific response takes place at the eastern edge of the warm-pool. These areas are examined in the three following subsections. 3.2. Kelvin Wave Forcing [11] The zonal current difference between the two experiments (Figure 3f ) shows an eastward anomaly that propagates from the WWE region to the eastern Pacific at a phase speed around 2.8 m s1(the mean phase speed in the
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Figure 2. (a) Time-longitude evolution of the zonal wind stress averaged over 1N – 1S. The thick line represents the eastern edge of the warm-pool simulated in the REF experiment (defined by sq = 22.2). Wind stress vector superimposed to its zonal component over the Pacific Ocean at (b) February 28th and (c) March 14th. Contour interval is 0.25 dyn cm2.
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Figure 3. Time-longitude evolution of surface temperature (upper), salinity (middle) and zonal current (bottom) averaged over 1N –1S. The left panels are for the REF experiment, and the right panels are for the difference between REF and NWE experiments. The contour interval is 0.5C for temperature, 0.2 psu for salinity and 0.2 m s1 for currents. The thick line represents the eastern edge of the warm-pool simulated in the REF experiment (defined by sq = 22.2). The dashed line represents the path of the Kelvin wave.
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model of the first baroclinic mode in the central Pacific) and reaches the eastern boundary at the beginning of May. Its maximum amplitude is 0.5 – 0.6 m s1 in the central Pacific and 0.4 m s1 in the eastern Pacific. This propagating signal is characteristic of a large downwelling Kelvin wave forced by the WWE. [12] The Kelvin wave signature can also be found in ADCP measurements. Figure 4 shows the time evolution of the observed zonal current profile at 140W over the top 200 m from the beginning of March to the end of May together with the same data from the REF and NWE experiments. The zonal current structure is remarkably similar in TAO data and the REF experiment. A maximum is found at depth, in the Equatorial Under Current (EUC) core (100 m deep). As the WWE occurs (mid-March), the zonal current intensifies throughout the vertical. It reaches its maximum values just at the beginning of April in the EUC core, a maximum that is slightly underestimated in the model by 0.1– 0.2 m s1. The absence of intensification in the NWE (Figure 4c) clearly indicates that the WWE is responsible for that strengthening through Kelvin wave propagation, and underlies the model capacity to simulate the subsurface zonal current impact of the Kelvin waves. [13] In the model, the dynamical effect of the Kelvin wave can be shown along the wave path. Figure 5 displays the evolution of zonal current profiles from the dateline in mid-March to the eastern Pacific in early May (i.e. along the dashed line in Figure 3). The existence of the Kelvin wave in REF induces a large eastward increase of zonal currents in the top 250 m. Surface currents which are weakly negative in NWE (around 0.1 m s1) (Figure 5b) become strongly positive when the Kelvin waves are generated (surface currents are greater than 0.4 m s1 in REF) (Figure 5a). Below, the passage of this Kelvin wave also intensifies the EUC core. It changes from 0.7 m s1 in NWE to 1.2 m s1 in REF. The zonal current difference along the Kelvin wave propagation reveals an eastward current anomaly down to 250 m (Figure 5c). The current anomaly has a complex vertical structure with strong amplitudes at the surface, a minimum above the EUC core, and a secondary maximum just below the EUC core. Maximum zonal current anomalies are found in the eastern part of the basin where they reach more than 0.6 m s1 at 100W at 70 m. This structure is in contradiction with that predicted by linear wave theory whereby there is a monotonic decrease in velocity amplitudes with depth in the top 1000 m. However, it is in good agreement with the observations of Johnson and McPhaden [1993] who isolate the Kelvin wave signal in zonal velocity time series of TAO moorings at 140W and 110W. Following their conclusions, it is the interaction between the mean vertical current system (EUC, SEC) and the Kelvin wave that may be held responsible for such a vertical structure. To evaluate the mechanisms at work in that complex vertical structure, the difference between the zonal momentum trends of REF and NWE experiments are computed along the Kelvin wave propagation. The major terms contributing to the zonal current anomaly are the zonal pressure gradient, the zonal advection and the vertical advection (Figure 6). The major terms contributing to the zonal current anomaly are the zonal pressure gradient, the zonal advection and the vertical advection. All other terms are negligible and will not be discussed.
Figure 4. Time evolution of the 5-day mean zonal velocity profiles at 0, 140W for: (a) TAO mooring, (b) REF and (c) NWE experiments. Contour interval is 20 cm s1.
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Figure 5. Evolution of zonal current profiles (averaged over 1N –1S) from the dateline in mid-March to the eastern Pacific in early May (i.e., along the dashed line in Figure 3) for (a) REF, (b) NWE and (c) their difference REF - NWE. Contour interval is 10 cm s1. [14] The zonal pressure gradient (Figure 6a) would be the only important component for a wave in an otherwise motionless ocean: it reflects the expected monotonic decrease of zonal current anomaly with depth in the nobackground zonal flow theory. However, as seen in Figure 6b, zonal advection is the major term contributing to the wave current maximum just above the EUC core. In fact,
this term reflects the fact that zonal current anomaly advects the background zonal flow existing just below the EUC core prior to the Kelvin wave passage (Figure 5b). This positive advective trend associated with the subsurface maximum of zonal current anomaly increases from west to east, and is maximum at 100W where the strongest mean negative zonal current horizontal gradients are observed just below
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Figure 6. Evolution of the difference between the zonal momentum trend of REF and NWE experiments (averaged over 1N–1S) from the dateline in mid-March to the eastern Pacific in early May (i.e., along the dashed line in Figure 3). (a) zonal pressure gradient trend, (b) zonal advection trend, (c) vertical advection trend. Contour interval is 0.15 m s1 month1. the EUC core. Zonal advection is also partially responsible for the anomaly minimum found above the core. Contrary to the positive zonal advection trend that is maximum in the eastern Pacific, the negative zonal advection trend is maximum in the central Pacific west of 160W where the strongest positive zonal gradient above the EUC is found.
[15] The vertical advection of zonal momentum (Figure 6c) has also a dipole structure in the vertical. It induces an acceleration at depth and a surface deceleration. While such a structure contributes to the complex vertical structure of zonal current differences, its maximum acceleration is significantly weaker and deeper than the one induced by
LENGAIGNE ET AL.: OCEAN REPONSE TO MARCH 1997 WWE
zonal advection. Therefore in the model, the zonal advection is mostly responsible for the departure from the linear theory, whereas Johnson and McPhaden [1993] attributed this departure to the vertical advection term. It is hard to point out what is responsible for the difference between these two conclusions. These differences may be due to differences in the mean flow conditions between the periods investigated in our study and the one of Johnson and McPhaden [1993]. It can also be attributed to either model biases or to difficulties inherent to vertical velocity evaluation for the data or to the fact that the theory that underpins the discussion in Johnson and McPhaden [1993] is that of linearized wave-mean flow interactions [McPhaden et al., 1986, 1987] while there may be important nonlinear wavewave interactions in the model. [16] Associated with the surface current anomaly (Figure 3f), SST anomalies are observed along the Kelvin wave propagation with a 10 day delay (Figure 3b). In the central Pacific, an anomalous warming (reaching more than 1C) is observed between 160W and 100W that persists over more than two months. Closer to the coast of South America, at 95W, a very localized anomalous cooling (reaching 0.75C) associated with the Kelvin wave propagation is observed. The final SST anomaly pattern observed along the Kelvin wave is a warming located at the eastern boundary (90W-80W) in agreement with observations [Boulanger et al., 1999]. [17] Temperature profiles along the path of the Kelvin wave for both experiments are presented in Figure 7 together with their difference. The main signal associated with the Kelvin wave propagation is a positive temperature anomaly propagating eastward in the thermocline core at the Kelvin wave speed (Figure 7c). It is the signature of a deepening along the Kelvin wave propagation (reaching 35 m at 170W) (Figures 7a and 7b). The anomaly is particularly intense in the western part of the basin where vertical temperature gradients are strong. East of 150W, its amplitude significantly decreases and the model underestimates the observed anomaly (not shown). This feature is related to the model thermocline that is known to be more diffuse than observed in the east [Vialard et al., 2001]. [18] In the central Pacific, the SST anomalies (Figure 3b) are mainly created by zonal advection of temperature (not shown), rather than by both horizontal components as in the study of Harrison and Giese [1988]. Indeed, in their study Harrison and Giese [1988] consider an annual mean forcing over which is superimposed a WWE, so that meridional advection contributes to the anomalies through the tropical instability waves. Here, these waves are almost nonexistent in April 1997 [Vialard et al., 2001] because of the trade wind relaxation that occurs in March in phase with the seasonal cycle. Thus zonal advection of SST is the dominant term controlling SST warming along the equator in the central Pacific in agreement with observations [Wang and McPhaden, 2001]. The amplitude of the warming due to an alteration of the instability field in the presence of equatorial Kelvin waves therefore clearly depends on the season during which the WWE occurs. [19] To summarize, the Kelvin wave activity forced by the March WWE has an important impact on the subsurface temperature and currents in the central and eastern Pacific where the wave-background flow interaction intensifies and
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significantly modifies both temperature and zonal current vertical profiles associated to the Kelvin wave. The surface and subsurface temperature warming associated with Kelvin wave activity certainly participates to the conditioning of the eastern Pacific prior to an El Nin˜o onset. 3.3. Local Response [20] By local response, we mean the response that occurs in the March WWE forced region which is approximately located between 135E and 165E. The presence of this WWE induces the development of southeastward surface currents under the wind event : the zonal current reaches more than 0.8 m s1 (Figure 3e) and the meridional current more than 0.3 m s1 (not shown) at the maximum of the March WWE. The TAO equatorial zonal current vertical structure at 147E in the top 200 meters is displayed in Figure 8a and compared to the structures simulated in REF (Figure 8b) and NWE (Figure 8c). The REF currents are in quite good agreement with observations. During the lifetime of the wind event, an intense eastward jet with maximum speeds of 0.8 m s1 at 30 m in REF and 0.6 m s1 in TAO data develops in the upper layer. Zonal currents are fairly homogeneous in the top 80 m (100 m in observations). In the model as in the observations, subsurface current anomalies tend to develop in a direction opposite to the positive local wind stress about 15 days after the wind-forcing. As suggested by Cronin et al. [2000], an analysis of the zonal momentum balance at this location confirms that these reversing jets are primarily due to the interplay between wind-forcing and the compensating pressure gradient. [21] The temperature section (Figure 8e) shows that the southeastward currents in the surface layer occurring in association with the wind event are associated with an intense cooling over the upper 80 meters, with temperatures dropping by 1.2C between the end of February and the end of the March. This cooling is also seen in TAO observations (Figure 8d). It occurs over a large region of the Pacific west of 160E between 4N and 4S. Other case studies of the evolution of observed local SSTs under WWE conditions in the western Pacific warm-pool confirm this cooling tendency [McPhaden et al., 1992; Delcroix et al., 1993; Cronin and McPhaden, 1997; Feng et al., 1998; Vecchi and Harrison, 1999]. During this particular westerly wind event, because the cooling does not happen in the NWE experiment (Figure 8f) where identical heat fluxes are applied, the cooling observed between the two experiments cannot be strongly controlled by decreased incoming shortwave and increased latent heat flux in contrast to what mostly happens under WWE conditions as described by Cronin and McPhaden [1997]. [22] In our study, this cooling is linked to both meridional and zonal advection processes caused by the WWE windforced currents. In fact, the strong southeastward currents in the surface layer associated with the March WWE induce an advection of cold surface water from (130E; 5N) to (145E; 0N) (Figure 9). From February 28th, as the WWE starts to develop (Figure 2b), strong southeastward currents generated by both local negative meridional winds and Ekman current convergence develop north of Papua New Guinea while strong westward currents are seen to be maintaining the warm-pool to the west. At that time, meridional advection tends to bring cold water from the north of Papua New
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Figure 7. Evolution of the temperature profiles (averaged over1N–1S) from the dateline in midMarch to the eastern Pacific in early May (i.e. along the dashed line in Figure 3). (a) REF, (b) NWE and (c) their difference REF - NWE. Contour interval is 1C. Guinea to the western part of the warm-pool (between 130E and 140E). By March 14th, as the WWE reaches its maximum (Figure 2c), the wind-forced currents have now developed over the whole warm-pool between 5N–5S and far western Pacific cold waters are zonally advected to the east cooling the warm-pool from 140E to 160E. The strong episodic role of both zonal and meridional advection of heat
in controlling SST under WWE conditions has already been noticed in some observational studies [Vecchi and Harrison, 2000; Cronin and McPhaden, 1997; Feng et al., 1998]. [23] The low temperature region located around 130E; 5N results from the existence of the Mindanao Current. Supplied by the bifurcation of the North Equatorial Current at the Philippine Coast, the Mindanao Current carries North
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Figure 8. Time evolution of the 5-day mean zonal velocity profiles (left panel) and temperature profiles (right panel) at 0, 147E for: TAO mooring (upper), REF experiment (middle) and NWE experiment (lower). Contour interval is 10 cm s1 for velocity and 0.5C for temperature.
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Figure 9. Currents vectors superimposed on sea surface temperature for the REF experiment at (a) February 28th and (b) March 14th. Contour interval is 0.25C. Pacific Water toward a confluence region with South Pacific Water near 5N [Nitani, 1972; Masumoto and Yamagata, 1991; Qiu and Lukas, 1996]. In the model, cooling in this region is induced both by meridional advection of cold water north of 5N and by Ekman pumping due to a positive zonal wind stress curl at this location. These cold waters are observed in the Reynolds SST data and SST seasonal variations in this region shows that cooling is maximum from December to March when WWE are the most intense. 3.4. Displacement of the Eastern Edge of the Warm and Fresh Pool [24] The largest simulated SST anomalies observed in response to the March 1997 WWE are located on the Eastern Edge of the Warm and Fresh-Pool (EEWFP) (Figure 3b) and reach more than 2C in mid-April. At the same location, an intense SSS signal (about 0.8 psu, Figure 3d) is also observed. These SST and SSS signals are displaced more or less at the same velocity (0.8 m s1) (Figures 3b– 3d). This fast propagation is related to a large eastward zonal current perturbation (0.8 –0.9 m s1) trapped in a 40m-deep surface layer and propagating from March 10th to the beginning of April at the same speed as the SST and SSS signals across the eastern edge of the warm-fresh pool (Figures 3e – 3f). This intense zonal current perturbation is stronger than the geostrophic zonal current anomaly of about 0.4 m s1 that would be produced in this region by Kelvin wave activity only (Figure 3f). Although these large
eastward zonal current anomalies are initiated when the Kelvin wave packet crosses the dateline where the salinity front is initially located (Figure 3c), the particularly fast simultaneous displacement of SST, SSS and zonal current anomalies must obey a specific mechanism. [25] In agreement with observations [Picaut et al., 1996], zonal advection is the major term responsible for the displacement of SST and SSS on the EEWFP (not shown). In order to investigate how the intense zonal currents are generated at the front, the contributions of each term of the zonal momentum equation at the surface are now examined along the equator, at the location of the thermohaline front. The major terms contributing to the zonal current variability at the first level of the model (5m) are presented in equation (2); all other terms being negligible. (a) is the zonal advection of zonal momentum, (b) is the zonal pressure gradient and (c) is the difference between the momentum flux entering the surface (zonal wind stress) and leaving at 10m depth (via the vertical diffusion term at 10m). @t u¼
u@ u |fflfflffl{zfflfflxffl}
1 @x p r0 |fflffl{zfflffl}
Zonal advection
Zonal pressure gradient
ðaÞ
ðbÞ
þ fðtx Þz¼0 ðKz @z uÞz¼10m g=10m
|fflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflffl{zfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflfflffl} StressVertical diffusion
ðcÞ ð2Þ
[26] These terms are presented in Figure 10 about the sq = 22.2 where the current intensification occurs (i.e., at the
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Figure 10. Surface trends of zonal momentum equation at the eastern edge of the warm-pool (defined by sq = 22.2) for the REF experiment. (left) (a) zonal advection (thick line), (b) zonal pressure gradient (dashed line), (c) vertical diffusion (including surface wind stress) (dashed and dotted line) and (a+b+c) time derivative of the zonal current (black thick line). (right) (b) zonal pressure gradient (dashed line), (c) vertical diffusion (dashed and dotted line) and (b+c) their sum (black thick line). eastern part of the thermohaline front). The mechanism leading to the fast displacement of the warm-pool is now discussed focusing on the initiation, development and termination phases. [27] Prior to the March 1997 WWE, the zonal pressure gradient at the front which tends to accelerate zonal currents to the east is most of the time counterbalanced by weak easterlies occurring over the front (Figure 2a). This tends to equilibrate the zonal momentum balance. Thus surface zonal currents are weak at the front. When the March 1997 WWE occurs, this equilibrium is broken. When the easterlies disappear and are replaced by westerlies over the front during the first 10 days of March (Figure 2c) (c) switches from negative to positive. Then, both the pressure gradient (b) and weak westerlies contributes to an acceleration of the current across the front. This results in a swift current intensification (Figure 3f). [28] During this initiation stage, intense eastward zonal currents have then been generated across the front whereas no current intensification occurred just east of the front. This results in a strong negative zonal gradient of zonal current at the eastern edge of the front (Figure 3e). In the development phase, the zonal gradient of zonal currents allow the zonal advection terms (a) to strongly develop (see the increases of zonal advection after March, 10th in Figure 10a) which accelerates the currents eastward. Therefore, zonal advection of zonal momentum significantly contributes to the intensification of the zonal current developing at the front at the end of the initiation stage. [29] While the surface eastward current is now developed, westerlies are observed to switch back to easterlies over the front, as seen on Figure 2a, acting to decelerate the current
(Figure 10a) via (c). However, this deceleration is nearly cancelled by (b), the zonal pressure gradient, as seen in Figure 10b. (b) and (c) are, from then on, almost always equilibrated, and the intense eastward current is maintained at the front. This shows that the zonal pressure gradient associated with the front is crucial to ensure the development phase. Without this pressure gradient, the initial acceleration would die quickly due to (c). [30] As the balance between (b) and (c) is restored after the initiation phase ((a) dominates) and as tracers are essentially advected, zonal currents, temperature and salinity verify the same type of equation during the development phase, that is @t X ¼ u@x X
where X is T, S or u. The fact that X = F (x-ut) is solution of this equation ensures the simultaneous displacement of SST, SSS and u at the same velocity u. [31] Finally, this process is maintained during about a month as long as (c) (a decelerating term) associated with weak westward winds is counterbalanced by the zonal pressure gradient at the front. Once westward winds in the central Pacific become sufficiently strong, that balance disappears and the decelerating term dominates, acting to weaken and eventually wipe out these intense eastward currents. [32] To confirm the existence of that mechanism in observations, we first focus on Reynolds SST data. As previously stated in section 2, these data are used in the heat flux restoring term. However, as the described displacement of the eastern edge is dominated by zonal advection in the
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Figure 11. Equator-time evolution of the sea surface temperature averaged over 1N–1S for (a) REF experiment and (b) Reynolds data. Contour interval is 0.5C. model as in the observations [Wang and McPhaden, 2001], it still makes sense to compare simulated and observed SST. The Reynolds SST observations show the existence of a rapid displacement of the EEWFP. In the model as in observations, this displacement is particularly intense in the second half of March where it reaches 0.9 m s1 in the REF experiment and 1.1 m s1 in Reynolds data (Figure 11). It corresponds to a warm-pool eastward displacement of about 2000 km in a month whereas it reaches 2300 km in the Reynolds SST data. The major difference between REF and observations is a more diffuse front west of the 28.5C isotherm in the observations. This might be due to the fact that Reynolds SSTs are too smooth for this relatively small and fast pattern. In addition, the poor satellite coverage in these areas where convection was intense during that time period is likely to contribute to a smoother depiction of the front in this product. In addition to comparing with the SST data, observed zonal currents at 170W were compared with REF (Figure 12). Both show a strong intensification of the currents in the surface layers from end of March to the beginning of April in phase with both the model and observations in phase with the displacement of the dynamical and thermohaline fronts. At a depth of 30 m, the zonal current reaches 70 cm/s in the TAO mooring at the end of March 1997, whereas it reaches the same maximum intensity a bit later in the model. The amplification of the surface zonal current at the SST and SSS fronts from the ADCP current meter, as well as the moving warm waters in Reynolds data tend to confirm that this mechanism occurs in the real ocean.
4. Impact of the Salinity on the Dynamics of the March 1997 WWE 4.1. Impact of Salinity in the Zonal Pressure Gradient [33] We have pointed out that the existence of the thermohaline front was a major factor entering in the
development and maintenance of the intense zonal currents generated at the front in response to the March 1997 WWE. In particular, the existence of a zonal pressure gradient associated with the salinity front at the eastern edge of the warm and fresh pool is of primary importance to the observed displacement. However, the exact contribution of temperature and salinity in the pressure gradient needs to be clarified. To further pursue this analysis, the role of salinity in the pressure gradient is now examined. [34] Computing the contribution of temperature and salinity to the density gradient proves that salinity and temperature have, at that time, an equal contribution to the density gradient at the front where the intense zonal currents are found (not shown). This important contribution of salinity in the model seems quite realistic, as a comparison of SSS at 180W between REF and TAO data during 1997 (Figure 13) indicates that the SSS temporal variation is reasonably well reproduced by the model (the SSS variation is around 0.7 psu in observations, while it reaches 0.9 psu in REF). This result suggests that salinity plays a major role in the zonal current intensification. To better understand this, a new sensitivity experiment (NWF as New Water Flux) is performed in which CMAP [Xie and Arkin, 1998] precipitation are used instead of the NCEP precipitation to form the fresh water flux forcing. [ 35 ] This new fresh water flux forcing noticeably increases the precipitation over the warm pool. It essentially induces a strengthening of the salinity front at the EEWFP as it can be seen in Figure 13 (the front now spans 1.5 psu). The simulated salinity front in NWF is unrealistic for this time period. However, such a strong salinity front may be observed during other time periods such as during October 1994 at the 165E TAO mooring. Changing the mean distribution of SSS via precipitation has a large impact on the amplitude of the studied mechanism (Figure 14b). As a matter of fact, one can see that zonal currents at the front are
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Figure 13. Time evolution of sea surface salinity at 0N, 170W for TAO mooring (thick line), REF experiment (dotted line) and NWF experiment (dashed line). increased by about 30% compared to REF (Figure 14a), the zonal gradient of salinity is strongly increased along the displacement of the EEWFP, which is also consequently increased. [36] To explain this stronger development of the currents at the front, we turn again to the terms of the momentum equation integrated at the front as in section 3.4. Compared to Figure 10, the same balance is holding on Figure 15. However, the pressure gradient (b) (see equation (1)) during the initiation phase is stronger than in Figure 10 and the nonlinear term (a) is also dramatically increased. As in section 3.4, the same mechanism holds to explain the strong increase of the zonal advection (a) that participates to the strong current acceleration. The tighter salinity front in NWF increases the pressure gradient as compared to REF. This allows the initial acceleration to be stronger at the front than in REF thus creating a stronger zonal gradient of currents at the front then enhancing the zonal advection of momentum. Once this acceleration is set, (b) mainly counterbalances (c) ((b)-(c) is even slightly positive during the development phase) and SST, SSS and u are simultaneously displaced. This sensitivity experiment emphasizes the crucial role played by the tightness of the salinity front in the onset and intensity of the previously described nonlinear mechanism.
Figure 12. Time evolution of the 5-day mean zonal velocity profiles at 0, 170W for (a) TAO mooring, (b) REF and (c) NWE experiments. Contour interval is 10 cm s1.
4.2. Impact of the Barrier Layer [37] In the western Pacific, salinity is also known to influence the ocean dynamics and thermodynamics [Lukas and Lindstrom, 1991; Vialard and Delecluse, 1998a, 1998b] via the formation of a barrier layer (BL) isolating the surface mixed-layer from subsurface cooler waters. During the March WWE, the ORCA model simulates the presence of a barrier layer (Figure 16). The local effect of the WWE is clearly identified: before the beginning of the WWE, a thick BL (50 m) is present in the fresh pool region west of the salinity front. The existence of such a barrier layer results from a strong downwelling near the salinity front created by
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Figure 15. Same as Figure 10 for the NWF experiment. convergence between central Pacific salty waters and western Pacific freshwaters [Vialard and Delecluse, 1998a, 1998b]. At the beginning of the WWE, the presence of this barrier layer prevents upward mixing of cold thermocline water. But, as zonal winds intensify, the very energetic turbulent mixing produced tends to mix the upper layer down to the thermocline, thus gradually destroying the BL in the region of the forcing (Figure 16). However, cooling by entrainment mixing is globally reduced by the existence of this BL in the region of the WWE. [38] Another barrier layer is also simulated at the eastern edge of the warm and fresh pool and contributes to the zonal current intensification as follows. As the zonal current amplitude increases in early March, the current anomalies remain trapped in a shallow layer (40 m) of warm and fresh waters. As the currents advect the thermohaline front, this shallow layer is advected over cooler and saltier waters of the central Pacific which are downwelled [Vialard and Delecluse, 1998b]. This situation creates a barrier layer below the shallow surface layer contributing to maintain the intensity of surface zonal currents. [39] To precisely understand the exact effect of this barrier layer, an experiment Without Barrier Layer (WBL) is performed using the REF momentum, heat and water forcings. This is done as by Vialard [2001] and Masson et al. [2002] by taking out the effect of salinity in the vertical diffusion coefficients used to calculate vertical diffusion of momentum and tracers [Blanke and Delecluse, 1993]. Hence, vertical mixing only depends on the vertical temperature structure, and no barrier layer can be formed. The results, displayed in Figure 14c, are in agreement with the previous description. First, zonal currents in the western
Pacific are weaker as the wind momentum is distributed over a deeper layer. Thus the zonal convergence and the thermohaline fronts are weaker. Second, as the mixed-layer at the eastern edge of the warm-pool is much deeper, there is no trapping of zonal momentum as in REF. Altogether, the mechanism described in section 3.4 does not develop in that simulation and the eastern edge of the warm-pool is advected only by geostrophic currents of the Kelvin waves and weak surface Ekman currents. Therefore the displacement of the EEWFP under the action of this current is considerably slower: while the front was displaced by about 2000 km in REF, it is now only displaced by 1200 km in WBL.
5. Conclusion [40] The onset of El Nin˜o conditions in the equatorial Pacific during 1997 has been strongly suggested to be associated to a series of westerly wind events (WWE) from late 1996 to mid-1997. While the first wind event observed near 150E in December 1996 – January 1997 has a weak effect on the central and eastern Pacific, the second one in March 1997 of longer duration, larger fetch and stronger intensity seems to strongly modify the equatorial ocean structure. To study the oceanic response to this strong March 1997 WWE, a series of simulations has been performed with the ORCA global OGCM [Madec and Imbard, 1996; Madec et al., 1998]. Our study reveals that the ocean has essentially three types of response to this wind event. [41] The first response is the generation of a Kelvin wave that efficiently deepens the thermocline across the basin,
Figure 14. (opposite) Time-longitude plots of sea surface salinity superimposed on zonal current (color) averaged over 1N – 1S for (a) REF, (b) NWF and (c) WBL experiments. Contour intervals are 0.25 psu for salinity. See color version of this figure at back of this issue.
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Figure 16. Time-longitude evolution of the barrier layer thickness averaged over 1N–1S for the REF experiment. Contour interval is 10 m.
with larger positive temperature anomalies west of 150W where the simulated temperature vertical gradients in the thermocline are stronger. While the theoretical Kelvin wave structure should reveal a monotonic decrease of zonal currents over the first 200 m, the model vertical structure in zonal currents shows a maximum at the surface, a minimum near the top of the thermocline and a second maximum in the core of (and below) the thermocline. This complex vertical structure is the result of a wave-mean flow interaction similar to what was observed by Johnson and McPhaden [1993]. However, contrary to Johnson and McPhaden [1993] who argued that vertical advection was the dominant mechanism generating the second maximum at depth, the model strongly suggests that zonal advection of zonal momentum is actually responsible for that second maximum. At the surface, the propagation of the Kelvin wave packet is associated with a significant warming (up to 1C) in the central Pacific due to zonal advection of background SST. A recent coupled study [Kessler and Kleeman, 2000] suggested that adding weak positive SST anomaly can produce an El Nin˜o hindcast about 50% stronger trough a trade winds reduction. Therefore the Kelvin wave impact on SST is also expected to reduce easterlies in the central Pacific. [42] The second effect associated to the strong March 1997 WWE was a local cooling occurring in the western Pacific (up to 1C between 130E and 160E) in quantitative agreement with observations (Figure 8). This strong
cooling was found to result from an advection of cold waters from a region located around 130E;5N by strong southeastward currents generated under the wind event. These cold waters actually result from Ekman pumping due to a positive zonal wind stress curl and from meridional advection of surface cold North Pacific waters by the Mindanao Current every year from December to March. Thus, each WWE that extends north of the equator over that region is likely to induce a strong cooling as the one observed during March 1997. [43] The last ocean response generated by the March 1997 WWE occurs at the eastern edge of the warm and fresh pool. It consists in a displacement of the eastern edge of the warm pool in March – April 1997 much faster (0.8 – 0.9 m s1) than the one that would have been induced by geostrophic zonal currents associated to the wind-forced downwelling Kelvin wave. Prior to the wind event, the zonal pressure gradient at the front is largely in equilibrium with the vertical diffusion trend (term (c) in equation 1 of section 2). As the wind reverses, the equilibrium is broken, surface currents accelerate and zonal advection of zonal momentum contributes to the local acceleration of surface zonal currents. Then, the zonal pressure gradient and vertical diffusion trends tend to cancel each other again, leaving the dynamical and thermohaline fronts to be described by a simple advection equation by zonal currents. Both the dynamical and thermohaline fronts then show a
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Figure 17. Observed equator-time evolution of sea surface temperature, Outgoing Longwave Radiation zonal wind stress averaged over 1N – 1S. Contour interval is 0.25C, 15 W m2, and 0.25 dyn cm2, respectively. coupled displacement at a velocity close to 0.8 m s1. However, as the westward wind over the front intensifies, the sum of the zonal pressure gradient and vertical diffusion of momentum is no longer balanced and the latter acts to dampen the frontal current magnitude. The eastward surge dies out after a displacement of 2000 km in a month. A comparison to observations suggests that this interaction, which is by far the largest signal in response to the March WWE, occurs in the real ocean. [44] The combination of a strong cooling in the western Pacific and the fast displacement of the eastern edge of the warm and fresh pool corresponds to a global eastward progression of the warm pool of about 2000 km in a month. Simultaneously, tropical deep convection (as measured by Outgoing Longwave Radiation [Shinoda et al., 1998]) and westerly winds displayed a similar eastward displacement (Figure 17). Such a coupled displacement of high SSTs, deep convection activity and westerly wind anomalies suggest an important role played by the specific characteristics (fetch, duration and intensity) of the March 1997 WWE. Indeed, the large zonal extension of that intense event contributed both to favor the equatorial advection of cold waters from 130E–5N and to force intense currents at the eastern of the warm and fresh pool, while its duration allowed the growth of a larger oceanic response. Hence the simultaneous cooling in the western Pacific and warming in the central Pacific were then likely to initiate the eastward displacement of convection and thus of westerly winds. The nonlinear mechanism observed at the eastern edge of the warm-pool contributed to that initiation by intensifying the local zonal current response and therefore by advecting the thermohaline front more quickly. Salinity through its contribution to the zonal pressure gradient at the front, the formation of barrier layers in the western Pacific and at the location of the front played a crucial role in that eastward
displacement. This result points to the strong need for a better knowledge of surface and subsurface salinity conditions and a better representation of water fluxes as they may be crucial to improve predictions of the onset of El Nin˜o events. Understanding the respective role of the March 1997 WWE and of the initial oceanic conditions (temperature, salinity,. . .) in the onset of the 1997 – 1998 El Nin˜o is now under investigation using coupled general circulation models.
[45] Acknowledgments. This work was supported by the Programme National d’Etude du Climat (PNEDC). Computations were carried out at the IDRIS/CNRS. We gratefully acknowledge the computing support of Maurice Imbard, Edmee Durand and Arnaud Jouzeau. The authors would like to thank the TAO operational group of NOAA/PMEL for providing TAO data. We also want to thank the reviewers for their fruitful comments. We thank Dr. Keith A. J. Rodgers for final comments and careful re-reading of the manuscript.
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J.-P. Boulanger, P. Delecluse, M. Lengaigne, G. Madec, S. Masson, and C. Menkes, Laboratoire d’Oce´anographie Dynamique et de Climatologie (CNRS/UPMC/IRD), Universite´ Pierre et Marie Curie, Tour 26, 4eme etage, Case 100, 4, Place Jussieu, 75252, Paris, France. (mllod@ipsl. jussieu.fr)
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Figure 14. (opposite) Time-longitude plots of sea surface salinity superimposed on zonal current (color) averaged over 1N – 1S for (a) REF, (b) NWF and (c) WBL experiments. Contour intervals are 0.25 psu for salinity.
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