Subtropical Air-Sea Interaction and Development of Central Pacific El Ni?o

XIE Ruihuang1), 2), HUANG Fei1), *, and REN Hongli2), 3)

?

Subtropical Air-Sea Interaction and Development of Central Pacific El Ni?o

XIE Ruihuang, HUANG Fei, and REN Hongli

1),,,266100,2),,,96822,3),,,100081,

The standard deviation of the central Pacific sea surface temperature anomaly (SSTA) during the period from October to February shows that the central Pacific SSTA variation is primarily due to the occurrence of the Central Pacific El Ni?o (CP-El Ni?o) and has a connection with the subtropical air-sea interaction in the northeastern Pacific. After removing the influence of the Eastern Pacific El Ni?o, an S-EOF analysis is conducted and the leading mode shows a clear seasonal SSTA evolving from the subtropical northeastern Pacific to the tropical central Pacific with a quasi-biennial period. The initial subtropical SSTA is generated by the wind speed decrease and surface heat flux increase due to a north Pacific anomalous cyclone. Such subtropical SSTA can further influence the establishment of the SSTA in the tropical central Pacificthe wind-evaporation-SST (WES) feedback. After established, the central equatorial Pacific SSTA can be strengthened by the zonal advective feedback and thermocline feedback, and develop into CP-El Ni?o. However, as the thermocline feedback increases the SSTA cooling after the mature phase, the heat flux loss and the reversed zonal advective feedback can cause the phase transition of CP-El Ni?o. Along with the wind stress variability, the recharge (discharge) process occurs in the central (eastern) equatorial Pacific and such a process causes the phase consistency between the thermocline depth and SST anomalies, which presents a contrast to the original recharge/discharge theory.

CP-El Ni?o; subtropical forcing; recharge/discharge process; phase consistency; thermocline depth

1 Introduction

Typical El Ni?o is characterized by broad anomalous warmer sea surface temperatures (SST) in the eastern-to-central tropical Pacific. The ‘standard’ El Ni?o defined by Rasmusson and Carpenter (1982) develops from the west coast of South America and propagates westward into the equatorial central Pacific. However, a new type of El Ni?o has been observed with sea surface temperature anomaly (SSTA) developed and maturely centered in the central equatorial Pacific. Different names, such as ‘dateline El Ni?o’ (Larkin and Harrison, 2005), ‘El Ni?o Modoki’ (Ashok, 2007), ‘Central Pacific ENSO’ (Kao and Yu, 2009), or ‘Warm Pool El Ni?o’ (Kug, 2009), have been employed for this type of El Ni?o. In this study, the name ‘Central Pacific El Ni?o (CP-El Ni?o)’ is used to distinguish it from the conventional Eastern Pacific El Ni?o (EP-El Ni?o) due to their apparently different warming centers.

Significant advances have been obtained in understand-ingthe nature of CP-El Ni?o. Yeh(2009) foresaw the increasing future occurrences of CP-El Ni?o under global warming. Lee and McPhaden (2010) reported the latest intensification of CP-El Ni?o since the 1990s. Kug(2009, 2010) pointed out that the zonal advective feedback is more important than the thermocline feedback in the development of CP-El Ni?o. In most of those studies, however, the authors stood on a zonal view, and the meridional development of CP-El Ni?o has rarely been investigated. Some years earlier, Wang and Picaut (2004) showed the meridional extension of the El Ni?o SSTA from the tropical to extra-tropical central Pacific on both hemispheres at the onset phase after the mid-1970s, which is seldom observed before. And recently, Yu and Kim (2011) examined the influence of the North Pacific sea level pressure (SLP) variations on the meridional development of CP-El Ni?o. Such influence can be explained by the ‘Seasonal Footprinting Mechanism (SFM)’ proposed by Vimont(2001, 2003) as it reveals strong correlation and physical linkage between the preceding winter north Pacific SLP and the succeeding summer-to-winter equatorial wind stress, and further implicates the burst of ENSO. Moreover, some other evidences also show that the subtropical air-sea interactions would lead to the onset of El Ni?o events (Anderson, 2003, 2004; Chang, 2007). All these previous studies indicate that both zonal and meridional analyses should be conducted to better understand CP-El Ni?o. Therefore, in this study, the SLP forcing from the extra-tropical to the equatorial central Pacific is re-examined during the developing stage of the CP-El Ni?o, and the meridional heat exchange (the recharge/discharge process between the equatorial and off-equatorial regions) is calculated.

This paper is organized as follows. Data and methodology are introduced in Section 2. Section 3 investigates the main characteristics of CP-El Ni?o and the corresponding air-sea interaction. In Section 4, the heat budget and heat exchange is examined. And Section 5 summarizes the main findings.

2 Data and Methodology

Three datasets are used in this study: the SST data are obtained from the Met Office Hadley Centre Sea Ice and Sea Surface Temperature dataset (HadISST) (Rayner, 2003), the surface wind stress and oceanic variables and surface heat flux from the ECMWF Ocean Analysis System (ORA-S3, Balmaseda, 2008), and the SLP from the National Centers for Environmental Prediction–National Center for Atmospheric Research (NCEP-NCAR) reanalysis (Kalnay, 1996). The period of 1960–2009 is selected for this analysis because of the availability of all the three datasets. All the data are pre-processed by removing the temporal variability shorter than 3 months, and the anomalous quantities are obtained by subtracting the monthly climatological mean for the period of 1980–2009. For a reminder of this study, the number ‘0’ indicates the developing year, ‘1’ denotes the decaying year; and the season of winter, spring, summer and autumn represent the DJF, MAM, JJA and SON mean, respectively.

Fig.1 shows the standard deviation of October(0)–February(1) mean SSTA during 1960–2009, and the mean SSTA composite of CP-El Ni?o and EP-El Ni?o, following the classification by Kug(2009). The EP-El Ni?o events occurred in 1972/73, 1976/77, 1982/83, and 1997/98, and the CP-El Ni?o events in 1977/78, 1990/91, 1994/95, 2002/03, 2004/05 and 2009/10. As shown in Fig.1a, three peak areas of the standard deviation (SD) appear and they are located in: the east part of the Ni?o-4 region (170?W–150?W, 5?S–5?N), the west part of the Ni?o-3 region (150?W–120?W, 5?S–5?N) and the east part of the Ni?o-3 region (120?W–90?W, 5?S–5?N, Ni?o-3R region, hereafter), respectively (Fig.1a). Interestingly, the SD peak area in the Ni?o-3R region corresponds to the warming core of the EP-El Ni?o (Fig.1b), while the SDs in the other parts of the Ni?o-3 and Ni?o-4 regions correspond to the warm core of the CP-El Ni?o and the extension of the EP-El Ni?o. Therefore, considering the propagation of the EP-El Ni?o, the Ni?o-3R region can be treated as the source of EP-El Ni?o.

In order to obtain the SST anomalies, a linear regression of the total SST anomalies to the Ni?o-3R index was first performed with the regression coefficient, then the SST anomalies associated with Ni?o-3R (EP-El Ni?o) were defined as. After subtractingfrom the total SST anomalies, the remainder is a quantity independent of the EP-El Ni?o. As shown in Fig.2a, the Ni?o-3R index has a high correlation with the Ni?o-3 index (=0.91). The Ni?o-3R index captures larger amplitude of SSTA than the Ni?o-3 index during the 1969/70, 1982/83 and 1997/98 EP-El Ni?o events, but smaller in the 1994/95, 2002/03, 2004/05 CP-El Ni?o events. As displayed in Fig.2b, the SSTA regression pattern of the Ni?o-3R index shows a typical SSTA distribution of the EP-El Ni?o, with the maximum center of action in the Ni?o-3R region extending from the west coast of the South America. Moreover, the Ni?o-3R index and the corresponding SST anomalies show comparable improvement of the Ni?o-CT index (Ren and Jin, 2011). The SD of the residual SSTA can be seen in Fig.2c, with large SD values locating in the equatorial central Pacific. It is noted that a subtropical area with large SD values is located near the coast of the North America, and it extends southwestwards and reaches the large SD area on the equator, indicative of the SST connection between these two regions.

Fig.1 Standard deviation of (a) October–February mean SSTA during 1960–2009, and (b) October–February mean SSTA composite of CP-El Ni?o (shaded in ℃) and EP-El Ni?o (contours in ℃). The classification of CP-El Ni?o and EP-El Ni?o follows Kug et al. (2009). The boxes from left to right denote Ni?o-4 region (solid black line), Ni?o-3.4 region (dotted purple line) and Ni?o-3 region (solid black line), respectively.

A combined regression S-EOF method, which is similar to that used by Kao and Yu (2009), is performed to the covariance matrix of residual SST anomalies to identify the leading structure of the CP-El Ni?o. The S-EOF (Season-reliant Empirical Orthogonal Function, Wang and An, 2005) is an effective method to examine the seasonal evolution of the El Ni?o phenomenon. To apply such a method, the monthly data need to be reorganized to obtain a four-season sequential spatial domain so that the spatial vectors of S-EOF modes have a scale of four while the time series varies yearly. The S-EOF results are sensitive to the selection of seasonal sequences. Based on sensitivity tests, the seasonal sequence starting from MAM(0) to DJF(1) was selected to better show the SSTA developing pattern of the CP-El Ni?o. After the leading mode is obtained, all other variables are monthly or seasonally regressed onto the principle component (PC) to examine the air-sea interaction during the development of the CP-El Ni?o.

Fig.2 (a) Ni?o-3R and Ni?o-3 indices during 1960–2009, (b) linear regression of SSTA onto the Ni?o-3R index, (c) standard deviation of the residual SSTA after the Ni?o-3R related SSTA are removed. The boxes from left to right denote Ni?o-4 region (solid black line), Ni?o-3.4 region (dotted purple line), and Ni?o-3 region (solid black line), respectively.

3 Air-Sea Interaction

3.1 S-EOF Results

The statistical results from the S-EOF of residual SSTA in the tropical Pacific (120?E–80?W, 30?S–30?N) are shown in Fig.3. The leading S-EOF mode accounts for 17% of the total variance in the residual SSTA. The spatial patterns clearly display the seasonal development of the CP-El Ni?o SSTA. The spring (MAM) pattern shows a horseshoe shape with strong positive SSTA extending from subtropical north and south American shore to equatorial central Pacific and weak out-of-phase SSTA in the far eastern Pacific (Fig.3a). This horseshoe-like pattern persists through summer (JJA). The hemispheric asymmetry is clearly shown by the enhanced SSTA in the north branch and the nearly unchanged SSTA in the south branch (Fig.3b). Although the positive SSTA distribution is still horseshoe-like in autumn (SON), the warming center shifts to the equatorial central Pacific (Fig.3c). The warm water is intensified in the boreal winter (Fig.3d). Because of the extension of its influence into the far eastern Pacific, the negative SSTA signals can be reversed there. Generally, the winter time SSTA pattern matches the mature phase of the CP-El Ni?o, which has been obtained by means of composite and regression in the above-mentioned studies. Throughout the four seasons, the most remarkable signals are the propagation of positive SST anomalies from north subtropical Pacific into the equatorial central Pacific in spring and summer, and the movement of the stronger and broader warming center to the eastern Pacific in autumn and winter. However, in the south branch of the horseshoe, moderate positive SSTA is maintained and no clear propagation is observed through seasons.

Fig.3 The leading S-EOF mode of seasonal mean residual SSTA in the tropical Pacific during 1960–2009. The fractional variance is 17%. The seasonal sequence starts from MAM(0) to DJF(1) as shown in (a) to (d).

Fig.4a shows the PC of the leading S-EOF and the November–February mean Ni?o-3R index, and the asymmetry between warm and cold events can be seen in the figure. The skewness of PC is ?0.79 while that of peak-time Ni?o-3R index is 1.57. The negative skewness of PC indicates that the centers of most La Ni?a events are located in the central Pacific, which results in stronger cold events than warm ones there (Kug, 2009). However, warm events occur more frequently than cold events, especially after 1985. The PC captures the historical CP-El Ni?o events of 1990/1991, 1991/1992, 1994/1995, 2004/2005, together with the strongest one of 2009/2010. Large amplitude is also shown for the 1986–1988 event due to its irregular development and statistical roughness in the S-EOF analysis. In the SSTA evolution map of this event, the SST anomalies first occur in the eastern Pacific in the summer of 1986, develop and mature in the same region in the spring of 1987, and then propagate into the central Pacific and stay there for the rest of 1987. The SST anomalies eventually retreat back to and vanish in the eastern Pacific in early 1988. These irregular SSTA footprints produce a large variance in the central Pacific, and the S-EOF method mathematically collects these variations and classifies them as a CP-El Ni?o event.

Fig.4 (a) normalized time series of the leading S-EOF mode (PC1, red bar) and November–February mean Ni?o-3R index (blue bar). The horizontal line denotes a one time standard deviation. (b) power spectra (thick blue line) of the PC1. The dashed black and red lines represent the 95% and 90% confidence level, respectively.

The leading periodicities of this type of SST variability are examined using the power spectral analysis (Fig.4b). The spectrum of PC shows that the period consists of a robust 2.5-year band and a statistically nonsignificant 10-year band, which indicates that the typical CP-El Ni?o has a decadal and quasi-biennial (QB) variability. The Pacific decadal variation period is not statistically significant; however, there is an evidence that the CP-El Ni?o is tightly linked to the decadal variation of the North Pacific Gyre Oscillation (Di Lorenzo, 2010) through the shallow subtropical overturning cells (STCs) (Kleeman, 1999; McPhaden and Zhang, 2002; Nonaka, 2002; Merryfield and Boer, 2005). Some studies also show that the subtropical decadal oscillations can affect the tropics through the change in surface wind stress (Barnett, 1999; Pierce, 2000; Karspeck and Cane, 2002; Liu and Alexander, 2007). Interestingly, the decadal oscillation is roughly corresponding to the interval of two large events. For example, the intervals between the 1986/87 and 1994/95, and between 1994/05 and 2004/05 warm events are approximately 10 years, as well as between the 1988/89 and 1998/99, and between the 1998/99 and 2008/09 cold events. On the other hand, the QB component well corresponds to the leading period of the conventional ENSO (Rasmusson and Carpenter, 1982; Barnett, 1991; Gu and Philander, 1995; Wang and An, 2005). New evidence also supports the fact that the QB oscillation does exist in the central Pacific (Liu, 2006; Bejarano and Jin, 2008; Yu, 2010). Because of the coexistence of the decadal and quasi-biennial variations in the CP-El Ni?o, the linear regression rather than composite is adopted in the following analysis in order to avoid inadequate sampling in a limited period.

3.2 Physical Interpretation

The results from section 3.1 present the seasonal SSTA evolution of the CP-El Ni?o. In this section, physical evidence is provided to illustrate how the subtropical air-sea interaction can produce local SSTA and induce the CP-El Ni?o. Sea level pressure (SLP), surface wind stress (TAU) and its zonal component (TAUX), sea surface temperature (SST), and net surface heat flux are projected onto the PC of the leading S-EOF in an expanded region (120?E–60?W, 30?S–60?N) covering the north Pacific. All the regression coefficients are plotted in Fig.5 with a 95% confidence level of statistical significance (two-tailed Student’stest). The seasonal footprint mechanism (SFM) is used to analyze physical processes and all the variables are displayed from the winter of year 0. If the north Pacific SLP anomalies are the first and foremost element for the generation of local SSTA, the SFM can be regarded as a causal chain initially occurring in winter (Vimont, 2003). The north Pacific SLP anomalies in Fig.5a (i) show a dipole SLP pattern with anomalous high and low pressure centers on each side of about 45?N, respectively. The area with negative anomalies extends southwestwards (Fig.5a (i)) and weakens the trade winds throughout the subtropics. The reduced trade winds are essential for the occurrence of anomalous local SSTA (Fig.5a (iii)), because the reduced wind speed decreases wind-induced-evaporation and convection, and, as a result, strong solar radiation would produce positive SSTA band extending towards the central tropical Pacific along with wind anomalies. This process is consistent with the wind-evaporation-SST (WES) mechanism (Xie and Philander, 1994) in both zonal and meridional directions. Note that the weak westerly anomalies in the western Pacific can generate downwelling oceanic Kelvin wave propagating eastward to deepen the thermocline and produce sub-layer warming in the central Pacific (Fig.5a (v)). Meanwhile, the anomalous easterlies appear in the tropical eastern Pacific due to the continuity of air flow in the subtropics. Furthermore, these easterly anomalies will strengthen the trade winds, induce thermocline shoaling, and produce the cooling in the eastern Pacific. In the spring(0) season, even though the SLP dipole pattern and southwesterly anomalies become weaker than those in winter(0), as suggested by Vimont(2001, 2003, 2009), the WES mechanism still serves as an positive feedback to maintain the subtropical SSTA band through the positive net heat flux (Fig.5b (iv)). The persistent subtropical SST anomalies then start to reinforce the westerly anomalies in the western and central Pacific in summer to initiate the equatorial SSTA by enhancing downward net sea surface flux (Fig.5c (iv)). It is important to notice that the north Pacific cyclonic anomaly in this season is replaced by an anomalous subtropical cyclone, which is the atmospheric response to the established equatorial SSTA, rather than a survival of one in winter. During spring(0) and summer(0), the anomalous westerlies extend eastward and the nodal point is located near 160?W, where the maximum subsurface temperature anomalies are found. In autumn(0), positive SST anomalies are established in the central tropical Pacific (Fig.5d (iii)), and trigger the cyclonic atmospheric response (Fig.5d (i)) in both hemispheres. These SST anomalies are intensified rapidly by the westerly induced Kelvin wave, and in boreal winter(1), the warm pattern of anomalous SST is well developed in the central equatorial Pacific and a ‘tail’ is extended into the eastern Pacific (Fig.5e (iii)). Also, the subsurface temperature anomalies expand eastward and occupy the thermocline layer in the eastern Pacific, which is also responsible for the ‘tail’ of the SST anomalies.

Fig.5 Seasonal evolution of linear regression coefficients of CP-El Ni?o. The first column (a–e i) shows the sea level pressure (unit in hPa, contour interval: 0.2) and surface wind stress (vector in dyn cm-2); the second column (a–e ii) denotes the zonal wind stress (unit in dyn cm-2, contour interval: 0.03); the third column (a–e iii) displays the SSTA (unit in ℃, contour interval: 0.2); the fourth column (a–e iv) represents the net downward surface heat flux (unit in Wm-2, contour interval: 3.0); the fifth column (a–e v) shows the subsurface seawater temperature (unit in ℃, contour interval: 0.2) and the depth of the thermocline (blue line). Note that the zero line in the surface heat flux is not drawn. The shadedareas indicate the correlation coefficients with a significant level of 95% (Student’s t test).

The above analyses present the evolving pattern of the CP-El Ni?o through the subtropics-to-tropics air-sea interaction. Importantly, the initial warming starts from the north subtropical Pacific and the central Pacific rather than from the far eastern equatorial Pacific. It is only in the mature phase can the positive SSTA stretch into the east Pacific. Since the the SST warming and atmospheric heating (net heat flux) are confined in the central Pacific, the induced Rossby and Kelvin waves produce the low-level westerlies and easterlies to the west and east of the warm SST region, respectively. Therefore, the easterly anomalies in the eastern tropical Pacific persist throughout the developing and mature phases, and the equatorial westerly anomalies are trapped in the region west of 160?W. Thus, two features of the CP-El Ni?o SSTA and wind stress can be described as: 1) the initial SSTA develops from the north subtropical and the equatorial central Pacific, and shows a out-of-phase structure in the central and eastern Pacific during the development but can only influence the eastern Pacific when it matures; 2) the westerly anomalies are confined in the west of 160?W without the clear signal of eastward propagation, while the easterly anomalies prevail in the eastern Pacific even in the mature phase. To summarize, the CP-El Ni?o initiated from the northeastern subtropics is a local atmosphere–ocean coupling phenomenon.

4 The Role of Atmospheric Forcing and Oceanic Advection

4.1 Mix Layer Heat Budget

In order to examine the relative importance of atmospheric forcing and oceanic advection and their contribution to the growth of the CP-El Ni?o, a mix layer heat budget analysis is performed. A transform version of the mix layer temperature equation (Kang, 2004) is derived as follows:

, (1)

where the variables with an overbar indicate the monthly climatology, and those without the overbar denote the monthly mean anomaly. The variablesandindicate the zonal and meridional currents averaged over the mix layer (top 50m), respectively.andindicate the oceanic temperature averaged over the sub-layer (50–100m) and the mix layer, respectively.wis the upwelling velocity at the bottom of the mix layer (50m), andis the depth of the mix layer, which is equal to 50m here .is the net heat flux,is the density of seawater, andCis the ocean heat capacity.

The term on the left hand side of Eq. (1) is the temperature tendency. The first, second, and third terms on the right-hand side of Eq. (1) are the zonal advective feedback(), the anomalous vertical advection by the anom-alous vertical current and the mean temperature gradient (), and the thermocline re-displacement associated with the climatological mean upwelling (),respectively. As pointed out by An and Jin (2001), the third term is also related to the so-called thermocline feedback. All three terms mainly serve as positive contribution to the growth of ENSO (Jin, 2006). The fourth and fifth terms are the anomalous meridional advection due to the mean and anomalous meridional current. The last term is the temperature tendency related to the surface heat flux. Other terms are neglected in the equation because of their relatively small values or damping effect of SST (Kug, 2009), which will not be shown in this study.

Fig.6 shows the seasonal evolution of the temperature tendency term averaged between 170?E and 120?W. Temperature tendency shows weak positive sign in subtropical regions during winter(0) and large values appear on the equator in spring(0) (Fig.6a). The warm tendency is contributed by(Fig.6d) and the surface heat flux (Fig.6e), which further demonstrates that the initial equatorial SSTA develops from the atmospheric heating induced by the anomalous north Pacific cyclone. It is noted that the late development of the zonal advective feedback (Fig.6b) and the thermocline re-displacement (Fig.6c) occur in summer(0) as the westerly anomalies are reinforced in the central Pacific. Driven by the westerly anomalies, the warm water transported from the warm pool increases the SST in the central and eastern Pacific; meanwhile associated with the eastward propagation of the oceanic Kelvin waves the mean upwelling is reduced and the thermocline layer is deepened. These two processes together result in the development of SSTA in the central Pacific and the first three terms on the right hand side of Eq. (1) are maintaining the warm tendency through summer(0) and autumn(0). However, as the SST increases, more convective clouds are produced by evaporation, which blocks solar radiation and cools down the SST. Consequently, the net heat flux turns negative and causes ocean cooling (Fig.6e). One of the meridional advection terms,(Fig.6f), tends to be in phase with the equatorial temperature tendency, while the other term,(Fig.6g), shows an out-of-phase tendency. Some early studies showed that the former term () is large and positive near the equator during the development of ENSO (Battisti, 1988; Lau, 1992;Kang, 2001). In this study, large positive contributions of this advection term have been calculated during autumn(0) as tropical SST anomalies were established. The mean poleward current transports warmer water away from the equator, and warms up off-equatorial regions. During the transition phase, the SST cooling is mostly due to the negative contribution of zonal advective feedback and the surface heat flux loss. The thermocline re-displacement and the anomalous meridional advection terms by meancurrent and anomalous temperature gradient (), however, do not contribute to the phase transition, but these two terms do accelerate the SST cooling tendency during the CP-El Ni?o decaying phase.

Fig.6 Latitude-time diagram of the regressed heat budget terms listed in Eq. (1). All terms are averagedbetween 170?E and 120?W. The unit for each term is ℃mon-1. Note the different contour intervals used for different terms. The shadedareas indicate the correlation coefficients with a significant level of 95% (Student’s t test).

4.2 Recharge/Discharge Process

According to the recharge oscillator theory proposed by Jin (1996, 1997a, b), the positive anomalies of the equatorial heat content or thermocline depth lead El Ni?o by a quarter of the ENSO cycle and the anomalous heat content is discharged after the mature phase of El Ni?o. This theory can well explain the life cycle of the EP-El Ni?o (Meinen and McPhaden, 2000; Burgers, 2005; Clarke, 2007). However, the study of the CP-El Ni?o requires more understanding of the heat content recharge/discharge (Kug, 2010; Horii, 2012). Here, using the anomalous wind stress curl, the recharge/discharge process is re-examined during the development of the CP-El Ni?o. The Sverdrup transport is calculated to evaluate the heat and mass exchange between off-equatorial and equatorial regions. Because the wind stress and anomaly in the tropical Pacific are dominated by the zonal component, the Sverdrup theory can be simplified in the following equation:

where theSvis the northward Sverdrup transport by wind driven currents, andτandτare the zonal and meridional wind stresses, respectively, andis the change rate of Coriolis parameter with latitude. Using Eq. (2), the Sverdrup transport above the thermocline is calculated and the averaged transport between 5?N and 15?N in the North Hemisphere and that between 5?S and 15?S in the South Hemisphere are shown in Fig.7. The positive value indicates the northward mass or heat transport and corresponds to discharge (recharge) process in the North (South) Hemisphere. Fig.7 shows that the Sverdrup transport in the North Hemisphere (NH) is much weaker than that in the South Hemisphere (SH). This hemisphere asymmetry was reported earlier in the heat content analysis of the warm pool El Ni?o composite completed by Kug(2009). Such an asymmetry is different from that of the EP-El Ni?o case, in which the Sverdrup transport in the NH is larger than that in the SH (Kug, 2003). It is also noted that the recharge and discharge processes occur at the same time in both hemispheres. The discharge (recharge) process takes place in the region east (west) of 150?W during almost the entire life cycle except summer(1) and autumn(1) in the South Hemisphere. The recharge (discharge) process in the central and eastern Pacific in the NH (SH) peaks in boreal spring(0), leading the CP-El Ni?o mature phase by three seasons, which seems to be in line with the recharge oscillator theory. However, rather than changing to a discharge process, the recharge process in the NH continues to evolve in the central and eastern Pacific after spring(0). In the SH, the recharge process decays after spring(0), but redevelops in autumn(0) and reaches a second peak in spring(1), and the discharge process occurs in the whole basin after autumn(1). Although the discharge process prevails and reaches the peak in the central and western Pacific in both NH and SH, the local SST warms up instead of cooling down during the developing phase, which is another proof of the fact that the wind-induced depth change of the thermocline layer is not effective in changing the SST in the central Pacific. In other words, the discharge/recharge process in the CP-El Ni?o is not completely consistent with the recharge oscillator theory.

Fig.7 (a) Longitude-time plot of the regressed Sverdrup transport averaged between 5?N and 15?N. The unit is kgsm, and the contour intervals are 10. (b) is same as (a), but for the transport averaged between 5?S and 15?S. The shadedareas indicate the correlation coefficients with a significant level of 95% (Student’s test).

Unlikethe case of EP-El Ni?o, the Sverdrup transport during the CP-El Ni?o shows a hemisphere asymmetry pattern with stronger transport occurring in the SH. It was pointed out in previous studies that the El Ni?o phase transition is accompanied by the southward shift of the westerly anomalies in the central equatorial Pacific, and such an shift can further cause the meridional asymmetry of the Sverdrup transport with a stronger transport in the NH than in the SH (Harrison, 1987; Harrison and Vecchi, 1999; Kug et al., 2003). Fig.8 shows the meridional distributions of the zonal wind stress anomalies averaged over the central Pacific (170?E–160?W) and the eastern Pacific (160?W–100?W). In the central Pacific, the westerly anomalies occur over the northern part of the tropical Pacific during the life cycle of the CP-El Ni?o (Fig.8a) because of the existence of the anomalous northern Pacific cyclone before autumn(0) and then the subtropical cyclones excited by central Pacific SSTA after autumn(0) (Fig.5). Therefore, the meridional gradient of wind stress and the Sverdrup transport are weak (Fig.8c). However, there are no such anomalous cyclones in the southern part of the central tropical Pacific, and the Rossby waves component of Gill-type atmospheric response (Gill, 1980) to the CP-El Ni?o SSTA is much weaker, and thus the wind stress curl and Sverdrup transport reach the peak at the zero wind isoline (Fig.8c). In the eastern Pacific, the weakening of the north Pacific cyclone anomaly decreases the negative wind stress curl and the southward Sverdrup transport in the northeastern part of the tropical Pacific during the development of the CP-El Ni?o. However, after summer(0), the Kelvin waves component of Gill-type atmospheric response generates negative wind stress curl, which causes the burst of the southward Sverdrup transport (Fig.8d). On the other hand, in the southern part of the eastern tropical Pacific, the existence of the westerly anomalies causes the persistent negative wind stress curl (Fig.8b) and produces the northward transport (Fig.8d). Thus, the hemispheric asymmetry of the anomalous west- erlies is the key reason for the hemispheric asymmetry of the Sverdrup transport.

Fig.9a shows the regressed monthly total Sverdrup transport into the region between 10?S and 10?N averaged over the Pacific basin. The transport has a positive sign during the developing phase of the CP-El Ni?o, which indicates an inrush of mass and heat into the equatorial region. The transport decreases nearly to zero around the peak stage of the CP-El Ni?o, reflecting a balance between the recharge and discharge process. The Sverdrup transport drops below zero in the decay stage as the influence of the discharge process exceeds that of the recharge process.

Fig.8 (a) Latitude-time diagram of the regressed zonal wind stress averaged in the central Pacific (170?E–160?W) and (b) the eastern Pacific (160?W–100?W). (c) and (d) are same as (a) and (b), respectively, but for Sverdrup transport. The units for zonal wind stress and Sverdrup transport are dyncm-2 and kgs-1m-1, and the contour intervals are 0.01 and 1011, respectively. The shadedareas indicate the correlation coefficients with a significant level of 95% (Student’s t test).

Fig.9 (a) Regressed and averaged monthly total Sverdrup transport in the Pacific basin between 10?S and 10?N. (b) same as (a), but for thermocline depth (depth of 20℃isotherm, Z20, blue line) and SSTA (red line).

Because the incoming Sverdrup transport is greater than the outgoing one, the basin mean thermocline depth would increase with an increasing rate of depth change before summer(0). The thermocline layer would get deeper until winter(0) as the incoming and outgoing Sverdrup transports reach the balance. The thermocline will become shallow after winter(0) during the decaying stage because the outgoing Sverdrup transport gradually exceeds the incoming one. This process of the thermocline depth change can be verified by the regressed basin mean thermocline depth and SST anomalies shown in Fig.9b. It is clear that the thermocline depth is in phase with the development of SSTA. Horii(2012) reported a similar relationship and declared that the interrelationship between the equatorial warm water volume and ENSO SSTA has changed from the quarter-cycle phase leading to phase consistency since 2000. This study of the Sverdrup transport indicates that the simultaneous occurrence of the recharge and discharge processes may be responsible for the phase consistence in the CP-El Ni?o.

5 Summary

The SSTA variability in the central Pacific has been studied by using an S-EOF analysis, and the atmospheric and oceanic thermodynamics have been examined based on the PC of the leading S-EOF mode. The standard deviation of the tropical Pacific SSTA during October and February shows that the SSTA variation in the central Pacific is due to the CP-El Ni?o events and has a close correlation with the northeastern subtropical SST variation. The leading S-EOF of the central Pacific SSTA clearly displays a seasonal equatorial SSTA evolving pattern associated with the subtropical variation from boreal spring to winter. The positive SSTA first appears in boreal spring covering the regions from the northeastern subtropical Pacific to the central equatorial Pacific, and it is maintained into summer. In autumn, the peak SSTA occurs in the central Pacific and continues to develop into a mature CP-El Ni?o in boreal winter. The spectral density of the PC indicates that the CP-El Ni?o has a robust quasi-biennial period and a statistically nonsignificant decadal period.

During the development of the CP-El Ni?o, the preceding winter anomalous North Pacific cyclone induces the anomalous southwesterlies, and decreases the trade winds and the evaporation in the subtropical eastern Pacific. As a result of increased solar radiation, the local SST is increased. This air-sea interaction becomes an initial element of the ‘Seasonal Footprint Mechanism’ (SFM), which continues to impact the central Pacific through the wind-evaporation-SST mechanism in the next several seasons. Before summer, the development of the subtropical SSTA band is mostly due to the positive surface heat flux anomalies into the ocean and the anomalous southwesterly-induced downwelling. The subtropical SSTA band reaches the equatorial central Pacific in summer, and begins to activate the anomalous zonal advective feedback and anomalous thermocline feedback, which are the main contributors to the growth of the equatorial SSTA from summer to winter. During the developing phase of the CP-El Ni?o, the equatorial westerly (easterly) wind stress anomalies are confined within the region west (east) of 160?W without obvious propagation. Controlled by such a wind pattern, the SST and subsurface temperature anomalies are also confined in the central Pacific except during the boreal winter, when the CP-El Ni?o SSTA can expand to further east of the Pacific. After the mature phase, the SST anomalies begin to show a cooling tendency because of the negative contribution of the anomalous zonal advective feedback and surface heat flux. The anomalous thermocline feedback and anomalous meridional advection by mean currents do not affect the phase transition but accelerate the decay of the CP-El Ni?o.

The Sverdrup transport is calculated to examine the heat and mass exchange through the recharge/discharge process between the equatorial and off-equatorial regions. The recharge and discharge processes occur simultaneously and display a hemisphere asymmetry pattern with a stronger Sverdrup transport in the SH. Such a hemispheric asymmetry is caused by the asymmetric distribution of the westerly anomalies in both the central and eastern tropical Pacific. Moreover, the simultaneous occurrence of the recharge and discharge processes may be responsible for the phase consistence between the thermocline depth and the SST anomalies.

In this study, the different stages of the CP-El Ni?o development are captured and analyzed. The results of the study are consistent with Yu’s (2010) and Yu and Kim’s (2011) work, and the findings of the hemispheric asymmetry supplement the conclusion by Kug(2003), and further verify and enrich the recharge/discharge theory.

Acknowledgements

Part of this work was completed while the authors were visiting the University of Hawaii at Manoa. We would like to thank Prof. Fei-fei Jin for his suggestions on this study. We are also grateful to the anonymous reviewers and editors for their suggestions and comments to improve this paper. Dr. Ruihuang Xie and Prof. Fei Huang were jointly supported by the National Basic Research Program of China (973 Program: 2012CB955604), National Natural Science Foundation of China (Nos. 40975038 and 40830106), and the CMA Program (GYHY200906008). Dr. Ruihuang Xie acknowledges the financial support provided by the China Scholarship Council. Dr. Hongli Ren is jointly supported by the 973 Program of China (2010CB950404), DOE grant DE-SC0005110, National Science Foundation (NSF) grants ATM 1034798, NOAA grand NA10OAR4310200.

An, S. I., and Jin F. F., 2001. Collective role of zonal advective and thermocline feedbacks in ENSO mode, 14: 3421-3432.

Anderson, B. T., 2003. Tropical Pacific sea-surface temperatures and preceding sea level pressure anomalies in the subtropical North Pacific., 108, 4732, DOI: 10.1029/2003JD003805.

Anderson, B. T., 2004. Investigation of a large-scale mode of ocean–atmosphere variability and its relation to tropical Pacific sea surface temperature anomalies., 17: 4089-4098.

Ashok, K., Behera, S., Rao, A. S., Weng, H., and Yamagata, T., 2007. El Ni?o Modoki and its teleconnection., 112: C11007, DOI: 10.1029/2006JC003798.

Balmaseda, M. A., Vidard, A., and Anderson, D. L. T., 2008. The ECMWF ocean analysis system: ORA-S3., 136: 3018-3034.

Barnett, T. P., 1991. The interaction of multiple time scales in the tropical climate system., 4: 269-285.

Barnett, T. P., Pierce, D. W., Latif, M., Dommenget, D., and Saravana, R., 1999. Interdecadal interactions between the tropics and the midlatitudes in the Pacific basin., 26: 615-618.

Battisti, D. S., 1988. Dynamics and thermodynamics of a warming event in a coupled tropical atmosphere–ocean model., 45: 2889-2919.

Bejarano, L., and Jin, F. F., 2008. Coexistence of equatorial coupled modes of ENSO., 21: 3051-3067, DOI: 10.1175/2007JCLI1679.1.

Burgers, G., Jin, F. F., and Oldenborgh, G. J., 2005. The simplest ENSO recharge oscillator.,32, L13706, DOI: 10.1029 /2005GL022951.

Chang, P., Zhang, L., Saravanan, R., Vimont, D. J., Chiang, J. C. H., Ji, L., Seidel, H., and Tippett, M. K., 2007. Pacific meridional mode and El Ni?o-Southern Oscillation., 34, L16608, DOI: 10.1029/2007GL030302.

Clarke, A. J., Gorder, V. S., and Colantuono, G., 2007. Wind stress curl and ENSO discharge/recharge in the Equatorial Pacific., 37: 1077-1091, DOI: 10.1175/JPO3035.1.

Di Lorenzo, E., Cobb, K. M., Furtado, J. C., Schneider, N., Anderson, B. T., Bracco, A., Alexander, M. A., and Vimont, D. J., 2010. Central Pacific El Ni?o and decadal climate change in the North Pacific Ocean., 3: 762-765, DOI: 10.1038/ngeo984.

Gill, A. E., 1980. Some simple solutions for heat induced tropical circulation., 106: 447-462.

Gu, D., and Philander, S. G. H., 1995. Secular changes of annual and interannual variability in the tropics during the past century., 8: 864-876.

Harrison, D. E., 1987. Monthly mean island surface winds in the central tropical Pacific and El Ni?o., 115: 3133-3145.

Harrison, D. E., and Vecchi, G. A., 1999. On the termination of El Ni?o., 26(11): 1593-1596, DOI: 10.1029/1999GL900316.

Horii, T., Ueki, I., and Hanawa, K., 2012. Breakdown of ENSO predictors in the 2000s: Decadal changes of recharge/discharge-SST phase relation and atmospheric intraseasonal forcing., 39, L10707, DOI: 10.1029/2012GL051740.

Jin, F. F., 1996. Tropical ocean-atmosphere interaction, the Pacific cold tongue, and the El Ni?o-Southern Oscillation., 274: 76-78.

Jin, F. F., 1997a. An equatorial ocean recharge paradigm for ENSO. Part I: Conceptual model., 54: 811-829.

Jin, F. F., 1997b. An equatorial ocean recharge paradigm for ENSO. Part II: A stripped-down coupled model., 54: 830-847.

Jin, F. F, Kim, S. T., and Bejarano, L., 2006. A coupled-stability index for ENSO., 33, L23708, DOI: 10.1029/2006GL027221.

Kalnay, E., Kanamitsu, M., Kistler, R., Collins, W., Deaven, D., Gandin, L., Iredell, M., Saha, S., White, G., Woollen, J., Zhu, Y., Leetmaa, A., Reynolds, B., Chelliah, M., Ebisuzaki, W., Higgins, W., Janowiak, J., Mo, K. C., Ropelewski, C., Wang, J., Jenne, R., and Joseph, D., 1996. The NCEP/NCAR 40-year reanalysis project., 77: 437-472, DOI: 10.1175/1520-0477(1996)077<0437:TNYRP>2.0.CO;2.

Kang, I. S., An, S. I., and Jin, F. F., 2001. A systematic approximation of the SST anomaly equation for ENSO., 79: 1-10.

Kang, I. S., Kug, J. S., An, S. I., and Jin, F. F., 2004. A near annual Pacific Basin mode., 17 (12): 2478-2488.

Kao, H. Y., and Yu, J. Y., 2009. Contrasting eastern-Pacific and central Pacific types of El Ni?o., 22: 615-632, DOI: 10.1175/2008JCLI2309.1.

Karspeck, A. R., and Cane, M. A., 2002. Tropical Pacific 1976–77 climate shift in a linear, wind-driven model.,32 (8): 2350-2360.

Kleeman, R., McCreary, J., and Klinger, B. A., 1999. A mechanism for generating ENSO decadal variability., 26(12): 1743-1746, DOI: 10.1029/1999GL900352.

Kug, J. S., Choi, J., An, S. I., Jin, F. F., and Wittenberg, A. T., 2010. Warm pool and cold tongue El Ni?o events as simulated by the GFDL 2.1 coupled GCM., 23: 1226-1239, DOI: 10.1175/2010JCLI3688.1.

Kug, J. S., Jin, F. F., and An, S. I., 2009. Two types of El Ni?o events: Cold tongue El Ni?o and warm pool El Ni?o., 22: 1499-1515, DOI: 10.1175/2008JCLI2624.1.

Kug, J. S., Kang, I. S., and An, S. I., 2003. Symmetric and antisymmetric mass exchanges between the equatorial and off-equatorial Pacific associated with ENSO., 108: 3284, DOI: 10.1029/2002JC001671.

Larkin, N. K., and Harrison, D. E., 2005. Global seasonal temperature and precipitation anomaly during El Ni?o autumn and winter., 32, L16705, DOI: 10.1029/2005GL022860.

Lau, N. C., Philander, S. G. H., and Nath, M. J., 1992. Simulation of ENSO-like phenomena with a low-resolution coupled GCM of the global ocean and atmosphere., 5: 284-307.

Lee, T., and McPhaden, M. J., 2010. Increasing intensity of El Ni?o in the central-equatorial Pacific., 37, L14603, DOI: 10.1029/2010GL044007.

Liu, Q. Y., Liu, Z. Y., and Pan, A. J., 2006. Conceptual model about the interaction between El Ni?o/Southern Oscillation and Quasi-Biennial Oscillation in far west equatorial Pacific., 49 (8): 889-896.

Liu, Z., and Alexander, M., 2007. Atmospheric bridge, oceanic tunnel, and global climatic teleconnections., 45, RG2005, DOI: 10.1029/2005RG000172.

McPhaden, M., and Zhang, D., 2002. Slowdown of the meridional overturning circulation in the upper Pacific Ocean., 415: 603-608.

Meinen, C. S., and McPhaden, M. J., 2000. Observations of warm water volume changes in the Equatorial Pacific and their relationship to El Ni?o and La Ni?a., 13: 3551-3559.

Merryfield, W., and Boer, G., 2005. Variability of upper Pacific Ocean overturning in a coupled climate model.,18: 666-683.

Nonaka, M., Xie, S. P., and McCreary, J., 2002. Decadal variations in the subtropical cells and equatorial Pacific SST.,29, 1116, DOI: 10.1029/2001GL013717.

Pierce, D. W., Barnett, T. P., and Latif, M., 2000. Connections between the Pacific Ocean Tropics and midlatitudes on decadal timescales., 13: 1173-1194.

Rasmusson, E. M., and Carpenter, T. H., 1982. Variations in tropical sea surface temperature and surface wind field associated with the Southern Oscillation/El Ni?o., 110: 354-384.

Rayner, N. A., Parker, D. E., Horton, E. B., Folland, C. K., Alexander, L. V., Rowell, D. P., Kent, E. C., and Kaplan, A., 2003. Global analyses of sea surface temperature, sea ice, and night marine air temperature since the late nineteenth century., 108, 4407, DOI: 10.1029/2002JD002670.

Ren, H. L., and Jin, F. F., 2011. Ni?o indices for two types of ENSO., 38, L04704, DOI: 10.1029/2010GL046031.

Vimont, D. J., Alexander, M., and Fontaine, A., 2009. Midlatitude excitation of tropical variability in the Pacific: The role of thermodynamic coupling and seasonality., 22 (3): 518-534, DOI: 10.1175/2008JCLI2220.1.

Vimont, D. J., Battisti, D. S., and Hirst, A. C., 2001. Footprinting: A seasonal connection between the tropics and mid-latitudes., 28 (20): 3923-3926, DOI: 10.1029/2001GL013435.

Vimont, D. J., Wallace, J. M., and Battisti, D. S., 2003. The seasonal footprinting mechanism in the Pacific: Implications for ENSO.,16: 2668-2675.

Wang, B., and An, S. I., 2005. A method for detecting season-dependent modes of climate variability: S-EOF analysis., 32, L15710, DOI: 10.1029/2005GL022709.

Wang, C., and Picaut, J., 2004. Understanding ENSO physics–A

review. In:Wang, C., eds., AGU Geophysical Monograph Series, 147: 21-48.

Xie, S. P., and Philander, S. G. H., 1994.A coupled ocean–atmosphere model of relevance to the ITCZ in the eastern Pacific., 46: 340-350, DOI: 10.1034/j.1600-0870.1994.t01-1-00001.x.

Yeh, S. W., Kug, J. S., Dewitte, B., Kwon, M. H., Kirtman, B., and Jin, F. F., 2009. El Ni?o in a changing climate., 461: 511-514, DOI: 10.1038/nature08316.

Yu, J. Y., and Kim, S. T., 2011. Relationships between extratropical sea level pressure variations and the Central Pacific and Eastern Pacific types of ENSO., 24: 708-720, DOI: 10.1175/2010JCLI3688.1.

Yu, J. Y., Kao, H. Y., and Lee, T., 2010. Subtropics-related interannual sea surface temperature variability in the central equatorial Pacific., 23: 2869-2884.

(Edited by Xie Jun)

10.1007/s11802-013-2143-7

ISSN 1672-5182, 2013 12 (2): 260-271

. Tel: 0086-532-66786326 E-mail: huangf@ouc.edu.cn

(August 29, 2012; revised October 30, 2012; accepted February 25, 2013)

? Ocean University of China, Science Press and Springer-Verlag Berlin Heidelberg 2013