1. Introduction
The western North Pacific anomalous anticyclone (WNPAC) is the most important and long-lasting, low-level anomalous circulation over the western Pacific warm pool on the interannual time scale, which has great impacts on both the monsoon and El Niño [see reviews by Wang and Li (2004), Li and Wang (2005), and Zhou et al. (2014) and the introduction of Wu et al. (2017, hereafter Part I)].
In Part I, we demonstrated that the maintenance of the WNPAC during El Niño mature winter and the following spring primarily relies on both remote forcing from the equatorial central-eastern Pacific (CEP) via the atmospheric bridge and the local air–sea interactions, with the former being dominant. The remote forcing works mainly via a moist teleconnection mechanism. The northerly component of the western flank of the northern branch of El Niño–related low-level twin cyclonic anomalies straddled along the equator over the CEP advects low moist enthalpy (dry) air into the tropical western North Pacific (WNP). The negative moist enthalpy advection suppresses convection over the tropical WNP and further stimulates the WNPAC. This is Part II of the study, in which we turn to the WNPAC formation mechanisms.
Figure 1 shows the temporal evolutions of the low-level circulation, precipitation, and SST anomalies from August (year 0) to December (year 0) month by month, which were obtained by regression onto the December (year 0)–February (year 1) [D(0)JF(1)] mean Niño-3.4 index (area-averaged SSTAs over 5°S–5°N, 120°–170°W). Hereafter, years 0 and 1 represent El Niño developing and decaying years, respectively. In August(0), the WNP is dominated by cyclonic anomalies (Fig. 1a). The cyclonic anomalies withdraw eastward in September(0) (Fig. 1c). In October(0), an anomalous anticyclone establishes over the South China Sea (Fig. 1e). The anomalous anticyclone moves into the Philippine Sea in November(0), along with the further eastward withdrawal of the cyclonic anomalies extending from the CEP (Fig. 1g), which represents the onset of the WNPAC. In December(0), the center of the anomalous anticyclone moves into the Philippine Sea (Fig. 1i), meaning that the WNPAC fully establishes. Corresponding to the eastward movement of the anomalous anticyclone from October(0) to December(0), the negative precipitation anomalies also extend eastward (Figs. 1e,g,i). In contrast, from August(0) to December(0), the sign of the SSTAs over the tropical WNP does not change (Figs. 1b,d,f,h,j). What processes cause the radical change in the low-level circulation anomalies over the tropical WNP under the circumstance that the underlying SSTAs change less?
(left) Monthly mean precipitation (shading; mm day−1) and 925-hPa streamfunction anomalies (contours; 106 m2 s−1) from August(0) to December(0) during the El Niño developing phase, which were obtained by regression against the D(0)JF(1)-mean Niño-3.4 index. The contour interval is 0.3. Solid (dashed) lines represent positive (negative) values. (right) As in (left), but for SST (shading; K) and 925-hPa wind anomalies (vectors; m s−1). The brown (green) lines denote the anomalous anticyclone (cyclone). For the precipitation and SST anomalies, only values reaching the 5% significance level are shown.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
(left) Monthly mean precipitation (shading; mm day−1) and 925-hPa streamfunction anomalies (contours; 106 m2 s−1) from August(0) to December(0) during the El Niño developing phase, which were obtained by regression against the D(0)JF(1)-mean Niño-3.4 index. The contour interval is 0.3. Solid (dashed) lines represent positive (negative) values. (right) As in (left), but for SST (shading; K) and 925-hPa wind anomalies (vectors; m s−1). The brown (green) lines denote the anomalous anticyclone (cyclone). For the precipitation and SST anomalies, only values reaching the 5% significance level are shown.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
(left) Monthly mean precipitation (shading; mm day−1) and 925-hPa streamfunction anomalies (contours; 106 m2 s−1) from August(0) to December(0) during the El Niño developing phase, which were obtained by regression against the D(0)JF(1)-mean Niño-3.4 index. The contour interval is 0.3. Solid (dashed) lines represent positive (negative) values. (right) As in (left), but for SST (shading; K) and 925-hPa wind anomalies (vectors; m s−1). The brown (green) lines denote the anomalous anticyclone (cyclone). For the precipitation and SST anomalies, only values reaching the 5% significance level are shown.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
Several mechanisms have been proposed to understand the formation of the WNPAC during the late fall of the El Niño developing phase. Wang and Zhang (2002) analyzed individual El Niño events and found that the formation of the WNPAC is very rapid and always concurs with an intraseasonal oscillation swinging from a cyclonic to an anticyclonic phase. Meanwhile, the enhanced cold air intrusion from midlatitude northeastern Asia due to a deepened East Asian trough during El Niño favors the establishment of the WNPAC. Lau and Nath (2006) demonstrated that these processes can be simulated by numerical experiments using a coupled global climate model (CGCM). This mechanism explains the abrupt establishment of the WNPAC. However, it should be noted that the WNPAC is an anomalous circulation on the interannual time scale. Hence, it should be regarded as a stationary atmospheric response to El Niño–related SSTAs, under some favorable background mean state. The physical processes, especially changes in the background mean state, that determine the onset timing of the WNPAC are not yet understood.
Some studies have proposed that the WNPAC shifts from South Asia (Chou 2004; Chen et al. 2007). The meridional component of the anomalous anticyclone transports dry and cold (moist and warm) air to the eastern (western) side of the anomalous anticyclone. The zonal asymmetry in the advection of the mean moist static energy (MSE) by anomalous wind tends to cause the anomalous anticyclone to shift eastward (Chou 2004). However, these studies did not explain what determines the shift speed of the anomalous anticyclone and why it moves into the tropical WNP in the late fall.
Recently, a new mechanism has been proposed by Stuecker et al. (2015). They argued that the formation of the WNPAC results from the interaction of the El Niño–Southern Oscillation (ENSO) cycle with the climatological annual cycle in the tropics through a mathematical power spectral analysis. However, detailed physical explanations about the formation of the WNPAC were not given.
In this study, we answer the following two unknown fundamental questions with respect to the formation of the WNPAC. First, what kind of changes in the background mean state has determined the generation of the WNPAC during the late fall of the El Niño developing phase instead of before? Second, which physical processes have linked the changes in the background mean state with the formation of the WNPAC?
The remaining sections of this part are arranged as follows. Section 2 introduces the observational and reanalysis datasets, analysis methods, and numerical models used in the study. Section 3 explores the formation mechanisms through observational analysis. Section 4 further investigates the related mechanisms through idealized numerical experiments. Finally, the major conclusions are summarized in section 5.
2. Datasets, methods, and models
a. Datasets
The datasets used in this part were 1) the Global Precipitation Climatology Project (GPCP) dataset (Adler et al. 2003), 2) the Met Office Hadley Centre Sea Ice and SST dataset HadISST; Rayner et al. 2003), and 3) the European Centre for Medium-Range Weather Forecasts (ECMWF) interim reanalysis (ERA-Interim; Dee et al. 2011). All these datasets cover the period 1979–2012.
b. Methods
c. Models
The state-of-the-art CGCM FGOALS-s2 (Bao et al. 2013; Zhou et al. 2014) and an anomaly atmospheric general circulation model (AGCM; Jiang and Li 2005) were used to conduct numerical experiments. A detailed introduction to the FGOALS-s2 model can be found in Part I. The anomaly AGCM was developed based on the Geophysical Fluid Dynamics Laboratory (GFDL) global spectral dry AGCM (Held and Suarez 1994). In this model, an idealized or realistic mean state can be specified so that either the evolution of an initial perturbation or the atmospheric response to a specific heating under a specified background mean state can be examined. For a detailed derivation of the governing equations, see Jiang and Li (2005). This model has been used to examine the effect of the winter and summer mean state on the modulation of the atmospheric response to ENSO forcing in the tropical Indo-Pacific region (Wang et al. 2003), the initiation of the boreal summer intraseasonal oscillation over the western Indian Ocean (Jiang and Li 2005), and the effect of background vertical shear on summertime synoptic-scale wave trains in the western North Pacific (Li 2006).
3. Observational analysis
Figure 1 shows that the formation of the WNPAC is associated with the following two processes. First, from October(0) to December(0), the negative precipitation anomalies gradually stretch eastward from the South China Sea (SCS) to the tropical WNP, corresponding to an eastward shift of the anomalous anticyclone (Figs. 1e,g,i). Second, the northern branch of the twin cyclonic anomalies driven by El Niño–related enhanced convection over the equatorial CEP gradually withdraws eastward from August(0) to December(0) (Figs. 1a,c,e,g,i), which leaves space for the formation of the WNPAC. The mechanisms responsible for the two processes were investigated as follows.
a. Establishment of the negative precipitation anomalies over the tropical WNP
In Part I, we demonstrated that the negative precipitation anomalies over the tropical WNP during El Niño mature winter are caused by anomalous descending motion, which is primarily driven by the negative moist enthalpy advection of northerly anomalies to the western flank of the northern branch of the twin cyclonic anomalies, namely, the negative 
(left) Monthly mean 925-hPa climatological specific humidity (shading; g kg−1) and anomalous wind from August(0) to December(0), which were obtained by regression against the D(0)JF(1)-mean Niño-3.4 index (vectors; m s−1). The meridional distribution of the meridional gradient of 925-hPa climatological specific humidity averaged over 125°–160°E is shown separately at the right of each panel (blue lines; 10−6 g kg−1). The area-averaged meridional gradient of 925-hPa climatological specific humidity over 1°–14°N, 125°–160°E is given at the top-right corner of each panel. (right) Corresponding horizontal advection of the climatological specific humidity by the anomalous wind, which is converted to energy units (W m−2). Values reaching the 5% significance level are stippled in white.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
(left) Monthly mean 925-hPa climatological specific humidity (shading; g kg−1) and anomalous wind from August(0) to December(0), which were obtained by regression against the D(0)JF(1)-mean Niño-3.4 index (vectors; m s−1). The meridional distribution of the meridional gradient of 925-hPa climatological specific humidity averaged over 125°–160°E is shown separately at the right of each panel (blue lines; 10−6 g kg−1). The area-averaged meridional gradient of 925-hPa climatological specific humidity over 1°–14°N, 125°–160°E is given at the top-right corner of each panel. (right) Corresponding horizontal advection of the climatological specific humidity by the anomalous wind, which is converted to energy units (W m−2). Values reaching the 5% significance level are stippled in white.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
(left) Monthly mean 925-hPa climatological specific humidity (shading; g kg−1) and anomalous wind from August(0) to December(0), which were obtained by regression against the D(0)JF(1)-mean Niño-3.4 index (vectors; m s−1). The meridional distribution of the meridional gradient of 925-hPa climatological specific humidity averaged over 125°–160°E is shown separately at the right of each panel (blue lines; 10−6 g kg−1). The area-averaged meridional gradient of 925-hPa climatological specific humidity over 1°–14°N, 125°–160°E is given at the top-right corner of each panel. (right) Corresponding horizontal advection of the climatological specific humidity by the anomalous wind, which is converted to energy units (W m−2). Values reaching the 5% significance level are stippled in white.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
The 
b. Eastward shift of northern branch of the twin cyclonic anomalies
The twin cyclonic anomalies straddled along the equator are Rossby wave responses to the positive heating anomalies over the equatorial CEP. Hence, the eastward withdrawal of the northern branch of the twin cyclonic anomalies from the tropical WNP to the central North Pacific should be caused, at least partially, by the eastward shift of the positive heating. The center of the positive heating anomalies moves eastward roughly 20° longitude from October(0) to December(0) (Figs. 1e,g,i).
As the underlying El Niño–related warm SSTAs remain nearly unchanged in shape and position from October(0) to December(0) (Figs. 1f,h,j), the eastward shift of the positive heating center cannot be attributed to changes in the lower boundary forcing. Our analysis suggests that the eastward shift is caused by the temporal evolution of the mean ITCZ through the following two mechanisms.
First, the climatological ascending motion around the date line intensifies remarkably from October to December, as the mean ITCZ retreats toward the equator. Meanwhile, the transition area between the climatological descending and ascending areas on the equator moves eastward (Figs. 3c–e). Climatological ascending (descending) motion tends to enhance (suppress) convection responses to the underlying SSTA forcing; that is, the region of climatological ascending motion has stronger responses to the underlying SSTAs than the region of climatological descending motion (Xiang et al. 2013a,b). Hence, the center of the positive precipitation anomalies gradually moves eastward from west of the date line to east (Figs. 3c–e).
Climatological 500-hPa ω (shading; 10−2 Pa s−1) and positive precipitation anomalies over the equatorial CEP (contours; mm day−1) from August(0) to December(0). The contour interval is 1.5 mm day−1.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
Climatological 500-hPa ω (shading; 10−2 Pa s−1) and positive precipitation anomalies over the equatorial CEP (contours; mm day−1) from August(0) to December(0). The contour interval is 1.5 mm day−1.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
Climatological 500-hPa ω (shading; 10−2 Pa s−1) and positive precipitation anomalies over the equatorial CEP (contours; mm day−1) from August(0) to December(0). The contour interval is 1.5 mm day−1.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
Second, the equatorial westerly anomalies to the west of the positive precipitation anomalies over the equatorial CEP advect climatological wet (high moist enthalpy) air eastward (positive
The eastward withdrawal of the northern branch of the twin cyclonic anomalies cannot be attributed to the eastward shift of the positive precipitation anomalies over the equatorial CEP entirely, especially during the summer and early fall. From August(0) to October(0), the center of the positive precipitation anomalies moves less, but the cyclonic anomalies withdraw eastward remarkably (Figs. 1a,c), suggesting that some other mechanism is at work. We further explore this issue in section 4b through an idealized numerical experiment using the dry anomaly AGCM.
4. Numerical experiments
a. Experiments using the CGCM FGOALS-s2
We conducted two pacemaker experiments by the FGOALS-s2, with upper-700-m ocean temperature in the equatorial CEP restored to the observational anomalies plus model climatology (red dashed triangles in Figs. 4 and 5). The two experiments were also used in Part I. In the first experiment, the atmospheric and oceanic components of the model are freely coupled outside the equatorial CEP (hereafter resCEP run), while in the second experiment, the upper-700-m ocean temperatures in ocean areas outside of the equatorial CEP are restored to the model climatology (hereafter resCEP_clmGLB run). We use the two experiments to further investigate whether the onset timing of the WNPAC can be reproduced realistically when the local air–sea interactions are suppressed. Furthermore, the mechanisms proposed from the observational analysis need to be verified by the numerical experiments. Brief descriptions of the two experiments are given in Table 1. A more detailed introduction can be found in Part I.
As in Fig. 1, but for the resCEP run. Red dashed triangles in the right panels denote the ocean areas, in which modeled upper-700-m ocean temperatures are restored to the observational anomalies plus the model climatology.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
As in Fig. 1, but for the resCEP run. Red dashed triangles in the right panels denote the ocean areas, in which modeled upper-700-m ocean temperatures are restored to the observational anomalies plus the model climatology.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
As in Fig. 1, but for the resCEP run. Red dashed triangles in the right panels denote the ocean areas, in which modeled upper-700-m ocean temperatures are restored to the observational anomalies plus the model climatology.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
As in Fig. 4, but for the resCEP_clmGLB run.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
As in Fig. 4, but for the resCEP_clmGLB run.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
As in Fig. 4, but for the resCEP_clmGLB run.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
Experiments using the FGOALS-s2 model.
In the resCEP run, the onset timing of the WNPAC was in November(0), which is consistent with that in the observational analysis (Figs. 1g and 4g). If we check closely the temporal evolution of the circulation anomalies over the tropical WNP, the simulation results only have some minor differences with the observation. In the observations, the cyclone anomalies over the tropical WNP withdraw eastward from August(0) to October(0) (Figs. 1a,c,e). An anomalous anticyclone establishes in the SCS (Fig. 1e). The eastern flank of the anomalous anticyclone stretches from the SCS into the Philippine Sea in November(0), which represents the onset of the WNPAC (Fig. 1g). In the following December(0), the center of the anticyclone stretches into the tropical WNP, which represents the full establishment of the WNPAC (Fig. 1i). The resCEP run reproduces the eastward withdrawal of the cyclone anomalies over the tropical WNP from August(0) to October(0) (Figs. 4a,c,e). However, the withdrawal speed is lower than that in the observation, so that the SCS is still covered by the cyclone anomalies (Fig. 4e). An anomalous anticyclone forms over the SCS and the Philippine Sea in November(0), with the center shifted northward relative to its counterpart in the observation (Figs. 1g and 4g). In December(0), the anomalous anticyclone extends farther eastward as in the observation (Figs. 1i and 4i). On the other hand, the formation processes of the cold SSTAs in the tropical WNP simulated by the resCEP run are also similar with that in the observation (Figs. 1 and 4). The cold SSTAs have fully established in September(0) via the air–sea interactions (Fig. 4d; Li et al. 2006), whereas the WNPAC cannot be established until November(0) (Fig. 4g).
In the resCEP_clmGLB run, the onset timing of the WNPAC was in December(0) (Fig. 5i), one month later than that in the observational analysis and the resCEP run. The experiment reproduces the eastward withdrawal of the cyclone anomalies from August(0) to October(0) (Figs. 5a,c,e). In November(0), the SCS and tropical WNP are still covered by very weak cyclone anomalies (Fig. 5g). The anomalous anticyclone forms over the SCS in December(0), with eastern flank stretching into the Philippine Sea (Fig. 5i). Because the SSTA development is largely suppressed, the cold SSTAs in the tropical WNP generated by the local air–sea interactions are far weaker than those in the resCEP run, especially in November(0) and December(0) (Figs. 5d,f,h,j). The difference in the onset timing of the WNPAC between the resCEP and resCEP_clmGLB runs indicates that the cold SSTAs in the tropical WNP associated with the local air–sea interactions make a significant contribution to the formation of the WNAPC in late fall.
In the observational analysis, the locations of the negative precipitation anomalies over the SCS and tropical WNP from October(0) to December(0) generally follow those of the negative 
As in Fig. 2, but for the resCEP run. The area-averaged meridional gradient of 925-hPa climatological specific humidity over 5°–15°N, 125°–165°E is given at the top-right corner of the left panels.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
As in Fig. 2, but for the resCEP run. The area-averaged meridional gradient of 925-hPa climatological specific humidity over 5°–15°N, 125°–165°E is given at the top-right corner of the left panels.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
As in Fig. 2, but for the resCEP run. The area-averaged meridional gradient of 925-hPa climatological specific humidity over 5°–15°N, 125°–165°E is given at the top-right corner of the left panels.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
The discrepancies of the resCEP run in simulating the negative 
b. Experiments using the dry anomaly model
From the observational analysis, we proposed that the eastward shift of El Niño–related positive heating anomalies over the equatorial CEP contributed to the eastward withdrawal of the northern branch of the twin cyclonic anomalies from the tropical WNP to the central North Pacific. However, there are two issues that have not been resolved. First, the change in the position of the positive heating anomalies cannot fully explain the eastward withdrawal of the cyclonic anomalies, especially from August(0) to October(0). During this stage, the central position of the positive heating anomalies moves eastward relatively slowly, while the western flank of the cyclonic anomalies withdraws from the SCS to the Philippine Sea (Figs. 1a,c,e). Second, in both the observations and the AGCM results, the dynamic processes are coupled with moist processes. Specifically, the cyclonic anomalies over the tropical WNP tend to enhance local convection through Ekman-pumping-induced boundary layer convergence in August(0) and September(0) (Figs. 1a,c). In turn, the stimulated positive precipitation anomalies over the tropical WNP may contribute to the westward stretch of the cyclonic anomalies. From October(0) to December(0), the positive precipitation anomalies disappear over the tropical WNP and the negative precipitation anomalies gradually stretch eastward from the SCS to the tropical WNP (Figs. 1e,g,i). It is not clear whether the eastward withdrawal of the cyclonic anomalies is a result or cause of the establishment of the negative precipitation anomalies over the tropical WNP and the WNPAC. These issues were resolved by idealized experiments using the dry anomaly model.
The dry anomaly model is driven by a specified three-dimensional heating field. In this study, we constructed the heating field using the horizontal pattern of the positive precipitation anomalies over the equatorial CEP during El Niño mature winter [D(0)JF(1) mean] (contours in Fig. 7), and the bow vertical profile with a maximum at about 300 hPa (figure not shown). The heating field was used to drive the model, with the background mean states specified as the climatology of August–December derived from the reanalysis data. The five sets of runs are referred to as HEAT-CEP-Aug, HEAT-CEP-Sep, and so on. The detailed experiment designs are listed in Table 2.
The 850-hPa streamfunction anomalies (shading; 106 m2 s−1) simulated by the (a) HEAT-CEP-Aug, (b) HEAT-CEP-Sep, (c) HEAT-CEP-Oct, (d) HEAT-CEP-Nov, and (e) HEAT-CEP-Dec runs. Contours are the horizontal distributions of the heating field used to drive the dry anomaly model (K day−1). The contour interval is 1 K day−1.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
The 850-hPa streamfunction anomalies (shading; 106 m2 s−1) simulated by the (a) HEAT-CEP-Aug, (b) HEAT-CEP-Sep, (c) HEAT-CEP-Oct, (d) HEAT-CEP-Nov, and (e) HEAT-CEP-Dec runs. Contours are the horizontal distributions of the heating field used to drive the dry anomaly model (K day−1). The contour interval is 1 K day−1.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
The 850-hPa streamfunction anomalies (shading; 106 m2 s−1) simulated by the (a) HEAT-CEP-Aug, (b) HEAT-CEP-Sep, (c) HEAT-CEP-Oct, (d) HEAT-CEP-Nov, and (e) HEAT-CEP-Dec runs. Contours are the horizontal distributions of the heating field used to drive the dry anomaly model (K day−1). The contour interval is 1 K day−1.
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
Experiments using the dry anomaly model.
In August and September, 
Climatological 850-hPa relative vorticity from August to December derived from the ERA-Interim dataset (10−6 s−1). The area-averaged meridional gradient of relative vorticity over 5°–20°N, 120°–160°E (black dashed boxes) is given at the top-right corner of each panel (s−1 m−1).
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
Climatological 850-hPa relative vorticity from August to December derived from the ERA-Interim dataset (10−6 s−1). The area-averaged meridional gradient of relative vorticity over 5°–20°N, 120°–160°E (black dashed boxes) is given at the top-right corner of each panel (s−1 m−1).
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
Climatological 850-hPa relative vorticity from August to December derived from the ERA-Interim dataset (10−6 s−1). The area-averaged meridional gradient of relative vorticity over 5°–20°N, 120°–160°E (black dashed boxes) is given at the top-right corner of each panel (s−1 m−1).
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
The above evidences indicate that even under the same forcing of the enhanced convection over the equatorial CEP, the stimulated cyclonic anomalies centered in the tropical central North Pacific during summer and early fall tend to be stretched westward into the tropical WNP, because of a stronger meridional gradient of climatological absolute vorticity over the tropical WNP. As a result, there is no space left for the formation of the WNPAC during this stage. In contrast, during late fall and winter, the cyclonic anomalies gradually withdraw eastward to a “proper position,” with the weakening of the meridional gradient of climatological absolute vorticity over the tropical WNP, so that the anomalous northerly component to western flank of the cyclonic anomalies can effectively suppress convection over the tropical WNP through negative moist enthalpy advection and further stimulate the WNPAC.
5. Summary
In this part, we focused on the WNPAC formation processes. During El Niño–developing summer and early fall, the tropical WNP is covered by the northern branch of the twin cyclonic anomalies straddled along the equator, which are stimulated by El Niño–related enhanced convective heating anomalies over the equatorial CEP. With the eastward withdrawal of the cyclonic anomalies, an anomalous anticyclone forms over the SCS in October(0), then gradually stretches into the Philippine Sea in November(0), and finally covers the tropical WNP in December(0). A schematic for the WNPAC formation mechanisms is shown in Fig. 9.
Schematic of the formation of the WNPAC. (a) During the early fall of El Niño developing phase [September(0)], the cyclonic anomalies stimulated by El Niño–related positive precipitation anomalies (green-filled ellipse) over the equatorial CEP extends westward to the SCS. The westward stretch of the cyclonic anomalies is enhanced by the positive meridional gradient of the mean relative vorticity (
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
Schematic of the formation of the WNPAC. (a) During the early fall of El Niño developing phase [September(0)], the cyclonic anomalies stimulated by El Niño–related positive precipitation anomalies (green-filled ellipse) over the equatorial CEP extends westward to the SCS. The westward stretch of the cyclonic anomalies is enhanced by the positive meridional gradient of the mean relative vorticity (
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
Schematic of the formation of the WNPAC. (a) During the early fall of El Niño developing phase [September(0)], the cyclonic anomalies stimulated by El Niño–related positive precipitation anomalies (green-filled ellipse) over the equatorial CEP extends westward to the SCS. The westward stretch of the cyclonic anomalies is enhanced by the positive meridional gradient of the mean relative vorticity (
Citation: Journal of Climate 30, 23; 10.1175/JCLI-D-16-0495.1
Two processes are essential for the formation of the WNPAC. The first is the eastward withdrawal of the cyclonic anomalies. As the cyclonic anomalies are a Rossby wave response to the enhanced convective heating over the equatorial CEP, the eastward shift of the positive heating anomalies from October(0) to December(0) contribute to the eastward withdrawal of the cyclonic anomalies. Furthermore, a more fundamental mechanism is revealed by an idealized numerical experiment using the dry anomaly model. The extent of the westward stretch of the cyclonic anomalies is modulated by the interactions with the mean state. As the Coriolis parameter increases poleward, the northerly anomalies to the western flank of the cyclonic anomalies tend to cause positive vorticity advection to the west and thus stretch cyclonic anomalies westward. From boreal summer to winter, the meridional gradient of the low-level climatological relative vorticity transforms from positive to negative over the tropical WNP, which causes the meridional gradient of the absolute vorticity to gradually decline. As a result, the westward stretch effect of the cyclonic anomalies gradually weakens, causing the eastward withdrawal of the cyclonic anomalies.
The second essential process is the formation of the negative precipitation anomalies over the tropical WNP, which drive the WNPAC to the northwest. As noted in Part I, the negative precipitation anomalies are primarily excited by the negative advection of the mean moist enthalpy by northerly anomalies to the western flank of the cyclonic anomalies excited by the enhanced convective heating over the equatorial CEP. From August(0) to December(0), the meridional gradient of the low-level climatological moist enthalpy over the tropical WNP, which is dominated by that of the climatological specific humidity, gradually transforms from positive to negative, because of the equatorward retreat of the ITCZ. Hence, only after November(0), when the tropical WNP is dominated by the negative meridional gradient of the climatological specific humidity, can the negative moist enthalpy advection anomalies establish there. This process determines the onset timing of the negative precipitation anomalies over the tropical WNP.
Except for above two processes, underlying cold SSTAs also modulate the onset timing of the WNPAC. In Part I, three pacemaker experiments by the CGCM FGOALS-s2 indicated that the WNPAC and associated local negative precipitation anomalies are largely maintained by the negative 
Finally, previous studies noted that the WNPAC formation is concurrent with early invasion of the East Asian cold air break into the Philippine Sea and the swing of the intraseasonal oscillation from a wet to a dry phase (Wang and Zhang 2002; Lau and Nath 2006; Li et al. 2016). In the study, we focused on the interannual variability based on monthly mean data and stressed the importance of the changes in the background mean states in the tropical WNP to the WNPAC formation. How the synoptic–intraseasonal variability modulates the formation of the WNPAC deserves further study.
Acknowledgments
This work is jointly supported by the China National Key R&D Program (2017YFA0603802), NSFC (Grants 41661144009, 41675089, and 41330423), 973 Project 2015CB453201, NSF AGS-1565653, and the Jiangsu Collaborative Innovation Center for Climate Change. This is SOEST publication number 10214, IPRC publication number 1283, and ESMC publication number 178.
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