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    Hovmöller diagrams of (a) OLR (W m−2), (b) TPW (mm), (c) JRA-25 ω300 (Pa s−1), and (d) JRA-25 ω850 (Pa s−1) during 1 Sep–30 Nov 2000. OLR and ω300 are averaged between 5° and 12.5°N, TPW are averaged between 5° and 7.5°N, and ω850 are averaged between 5° and 10°N. Black crosses indicate the reference days (DAY0s) in the composite analysis over the EP. In (a), contours indicate MRG wave–filtered OLR anomalies, which are extracted in a similar way to Wheeler and Kiladis (1999), except that an antisymmetric–symmetric decomposition is not used. Contours are drawn at −10, −5, 5, and 10 W m−2, and negative values are contoured in dotted lines. The white slanted lines in (b) are missing data.

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    Power spectra of (a),(b) OLR (W2 m−4); (c),(d) TPW (mm2); (e),(f) ω300 (10−3 Pa2 s−2); and (g),(h) ω850 (10−3 Pa2 s−2) in the boreal autumn during 2000–05. (a),(c),(e),(g) The WP (7.5°N, 150°E) and (b),(d),(f),(h) the EP (7.5°N, 130°W). In (e)–(h), black and gray lines indicate JRA-25/JCDAS and ERA-Interim data, respectively. Abscissa indicates periods (day) in the logarithmic scale with 91 days on the extreme left and 0.5 day on the extreme right. Ordinate indicates power spectra multiplied by frequency. Error bars indicate the confidence intervals at the 95% significance level.

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    Spectral power of (a) OLR (W2 m−4), (b) TPW (mm2), (c) ω300 (10−3 Pa2 s−2), and (d) ω850 (10−3 Pa2 s−2) integrated from the period of 3.03–7.00 days in the boreal autumn during 2000–05.

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    Composites of JRA-25/JCDAS BPF-winds (m s−1) at (a) 850 and (b) 1000 hPa over the EP at DAY0 in the boreal autumn during 2000–05. (c) As in (a), but for QuikSCAT wind anomalies (m s−1) at 10 m over the ocean. The composites are based on all 83 cases. Arrows represent composite winds where either the zonal or meridional component is significant at the 95% level.

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    Composite longitude–pressure cross sections of BPF-meridional winds (m s−1) at 850 hPa over the EP at DAY0 in the boreal autumn during 2000–05. Cross sections at (a) 7.5°N and (b) the equator are shown. The composite is performed for all 83 cases. Shades indicate significant regions at the 95% level.

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    Composite Hovmöller diagrams of (a),(c) JRA-25/JCDAS BPF-meridional winds (m s−1) at 850 hPa and (b),(d) QuikSCAT meridional wind anomalies (m s−1) at 10 m over the ocean in the boreal autumn during 2000–05. Cross sections are at (a),(b) 7.5°N and (c),(d) the equator. The composite is performed for all 83 cases. In (a) and (c), the contour interval is 0.2 m s−1, and values greater than or equal to 0 m s−1 are contoured in solid lines. In (b), contours are set at −1, −0.5, 0.5, and 1 m s−1. In (d), contours are drawn at −0.4, −0.2, 0.2, and 0.4 m s−1. Shades indicate significant regions at the 95% level. Thick dashed lines in (a) and (c) indicate the phase speeds of −9 and −20 m s−1, respectively.

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    As in Fig. 6c, but for a composite from DAY-7 to DAY7 for the 21 cases with the coupled disturbance of a vortex and a cross-equatorial circulation. The contour interval is 0.4 m s−1 and values greater than or equal to 0 m s−1 are contoured in solid lines. Thick solid and dashed lines indicate the phase speed of 20 m s−1 and the group velocity of 5 m s−1, respectively.

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    Composite longitude–pressure cross sections of (a),(b) BPF-divergence (s−1) and (c),(d) BPF-vertical pressure velocities (Pa s−1) at DAY0 in the boreal autumn during 2000–05. (a),(c) For the EP and (b),(d) for the WP. The composite is performed for all cases. In (a) and (b), contours are set at −5, −3, −2, −1, 0, 1, 2, 3, 5, 10, and 15 × 10−6 s−1, and values greater than or equal to 0 s−1 are contoured in solid lines. In (c) and (d), contours are set at −0.14, −0.1, −0.06, −0.02, 0, and 0.02 Pa s−1, and values greater than or equal to 0 Pa s−1 are contoured in solid lines. Shades indicate significant regions at the 95% level.

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    Composite time series of TRMM Q1QR profiles (a) over the EP (6.25°–8.75°N, 131.25°–128.75°W) and (b) over the WP (6.25°–8.75°N, 148.75°–151.25°E) in the boreal autumn during 1998–2007. The composite is performed for all cases. Contours are set at 1, 1.3, 1.6, 1.9, 2.2, 2.5, 3, 5, and 7 K day−1.

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    Composite vertical profiles of each term (10−5 m2 s−3) on the right-hand side of the budget equation of the eddy kinetic energy in Eq. (1). Profiles are averaged over (a) the EP (5°–15°N, 140°–120°W) and (b) the WP (5°–15°N, 140°–160°E). Black solid, black dashed, gray solid, and gray dashed lines indicate the conversion from the kinetic energy of the 3-month-averaged field to the eddy kinetic energy (KmKe), the conversion from the eddy available potential energy to the eddy kinetic energy (AeKe), the convergence of the geopotential fluxes (GKe), and the advection of the eddy energy by the mean flow (advKe), respectively. The composite is performed for all cases.

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    (a) Composite BPF-ω850 (black contours) and BPF-ω300 (blue contours) of the coupled disturbance over the EP at DAY0 in the boreal autumn during 1998–2007. The contours start from −0.04 Pa s−1 with an interval of −0.04 Pa s−1. Shades indicate the ratios of composite BPF-ω850 to BPF-ω300. Shades are plotted where both BPF-ω300 and BPF-ω850 are significant with the 95% level. Arrows indicate horizontal winds (m s−1) at 850 hPa where either the zonal or meridional component is significant at the 95% level. (b) Profiles of TRMM Q1QR (K day−1). Gray and black lines are for the profiles averaged over the southwestern part (6.25°–7.5°N, 135°–130°W) and the northeastern part (7.5°–8.75°N, 131.25°–125°W, and 7.5°–10°N, 123.75°W, and 8.75°–10°N, 122.5°–121.25°W) of the vortex, respectively. These analysis regions are indicated with gray boxes in (a). Error bars indicate confidence intervals with the 95% significance level. (c) As in (b), but for vertical pressure velocity profiles. The composite is performed for 35 cases with the coupled disturbance.

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    As in Fig. 8a, but for a composite of 35 cases with the coupled disturbance over the EP in the boreal autumn during 1998–2007. The contour interval is 1 × 10−6 s−1 and values greater than or equal to 0 s−1 are contoured in solid lines. The cross section is averaged between 6.25° and 10°N.

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    Composite latitude–pressure cross sections of BPF-divergence (shades; 10−6 s−1) over the EP at DAY0 in the boreal autumn during 1998–2007. Cross sections are averaged over (a) 140°–135°W, (b) 135°–130°W, (c) 130°–125°W, and (3) 125°–120°W. Purple contours indicate significant regions at the 95% level. Arrows indicate meridional and vertical wind velocity, and black arrows indicate vectors where either the zonal or meridional component is significant at the 95% level. Vertical velocities are stretched to appear clearly in the figures. The composite is performed for 35 cases with the coupled disturbance.

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    As in Fig. 12, but for composites of the cases with (a) the isolated vortex-type disturbances and the cases with (b) the isolated MRG wave–type disturbances.

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    Schematic for the relationships among cumulus convection, the synoptic-scale coupled disturbances, and the large-scale environment over the EP in the boreal autumn. Small and large circles indicate a vortex disturbance and an MRG wave–type disturbance, respectively. The thick light gray arrow represents deep convergence associated with the MRG wave–type disturbance. Dark gray arrows indicate shallow convergence, which is largely driven by the strong SST gradient. Black solid and open arrows indicate phase and energy propagation, respectively. Small and large cloud-shaped figures are depicted as congestus and well organized systems, respectively.

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Relationships between Rain Characteristics and Environment. Part II: Atmospheric Disturbances Associated with Shallow Convection over the Eastern Tropical Pacific

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  • 1 Atmosphere and Ocean Research Institute, University of Tokyo, Kashiwa, Chiba, Japan
  • | 2 Atmosphere and Ocean Research Institute, University of Tokyo, Kashiwa, Chiba, and Research Institute for Global Change, Japan Agency for Marine-Earth Science and Technology, Yokosuka, Kanagawa, Japan
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Abstract

Synoptic-scale westward-propagating disturbances over the eastern Pacific (EP) are analyzed in boreal autumn, utilizing spectral analysis, composite analysis, and energy budget analysis. The results are compared with those over the western Pacific (WP).

Spectral peaks of total precipitable water (TPW) and vertical velocity at 850 hPa (ω850), and outgoing longwave radiation (OLR) are detected at periods of ~3–7 days over the EP. Meanwhile over the WP, a spectral peak of OLR is pronounced, but peaks of TPW and ω850 are not detected. Composite analysis reveals that disturbances that have a coupled structure, with a vortex at its center near ~9°N and a mixed Rossby–gravity (MRG) wave–type disturbance, frequently exist over the EP. At the same time, the disturbances have a double-deck structure associated with divergence both in the upper and in the middle to lower troposphere. These disturbances are associated with both deep convection and congestus, which generate kinetic energy of the disturbance in the upper and in the lower troposphere, respectively.

Examining diabatic heating in relation to the coupled disturbances, deep heating with the peak at the height of ~7.5 km is greatest in the northeastern part of the vortex. The coupled MRG wave–type disturbance provides a relatively deep cross-equatorial southerly flow into the northeastern part of the vortex. It is suggested that deep rain is maintained with the existence of deep convergence produced by the coupled disturbances over the EP, where a very shallow convergence field exists on average.

Corresponding author address: Chie Yokoyama, Department of Atmospheric Sciences, University of Utah, 135 South 1460 East, Salt Lake City, UT 84112-0110. E-mail: chie.yokoyama@utah.edu

Abstract

Synoptic-scale westward-propagating disturbances over the eastern Pacific (EP) are analyzed in boreal autumn, utilizing spectral analysis, composite analysis, and energy budget analysis. The results are compared with those over the western Pacific (WP).

Spectral peaks of total precipitable water (TPW) and vertical velocity at 850 hPa (ω850), and outgoing longwave radiation (OLR) are detected at periods of ~3–7 days over the EP. Meanwhile over the WP, a spectral peak of OLR is pronounced, but peaks of TPW and ω850 are not detected. Composite analysis reveals that disturbances that have a coupled structure, with a vortex at its center near ~9°N and a mixed Rossby–gravity (MRG) wave–type disturbance, frequently exist over the EP. At the same time, the disturbances have a double-deck structure associated with divergence both in the upper and in the middle to lower troposphere. These disturbances are associated with both deep convection and congestus, which generate kinetic energy of the disturbance in the upper and in the lower troposphere, respectively.

Examining diabatic heating in relation to the coupled disturbances, deep heating with the peak at the height of ~7.5 km is greatest in the northeastern part of the vortex. The coupled MRG wave–type disturbance provides a relatively deep cross-equatorial southerly flow into the northeastern part of the vortex. It is suggested that deep rain is maintained with the existence of deep convergence produced by the coupled disturbances over the EP, where a very shallow convergence field exists on average.

Corresponding author address: Chie Yokoyama, Department of Atmospheric Sciences, University of Utah, 135 South 1460 East, Salt Lake City, UT 84112-0110. E-mail: chie.yokoyama@utah.edu

1. Introduction

In the easterlies, there are westward-propagating disturbances so-called easterly waves with periods of ~3–5 days and zonal wavelengths of ~2000–5000 km in the lower troposphere (Riehl 1945, 1954; Palmer 1952). These disturbances are known to be coupled with deep cumulus convection (e.g., Chang 1970). Such disturbances are observed in different tropical regions including the intertropical convergence zone (ITCZ) over the eastern Pacific (EP) and the warm pool region over the western Pacific (WP). Some of these westward-propagating as well as eastward-propagating disturbances are identified as “equatorial wave modes coupled with cumulus convection.” Various equatorial wave modes such as Kelvin wave, mixed Rossby–gravity (MRG) wave, equatorial Rossby wave, and inertial gravity wave with a common equivalent depth are identified, using the cloud data observed from satellites (Takayabu 1994; Wheeler and Kiladis 1999). Many studies have examined what kind of structure these convectively coupled equatorial waves have, and where they are dominant (e.g., Hendon and Liebmann 1991; Wheeler et al. 2000; Kiladis et al. 2009).

Over the tropical Pacific, westward-propagating synoptic-scale disturbances are active from the boreal summer to the autumn (Gu and Zhang 2001). Such westward-propagating signals in outgoing longwave radiation (OLR) are more dominant in the Eastern Hemisphere than the Western Hemisphere (Roundy and Frank 2004). As a result, more studies are found to depict the characteristics of disturbances over the central-western Pacific, where stronger signals in OLR and infrared (IR) brightness temperature are found compared to those over the EP. Takayabu and Nitta (1993) showed the existence of different disturbance types near 150°E and near the date line in the boreal summer. Near 150°E, deep vertical disturbances called tropical depression (TD) type are dominant. On the other hand, the MRG wave–type disturbances are dominant around the date line (Liebmann and Hendon 1990; Takayabu and Nitta 1993). Both types are coupled with deep convection.

Over the EP, some studies show different aspects of synoptic-scale disturbances. It is shown that there are MRG wave–type disturbances associated with deep convection over the central-eastern Pacific (Liebmann and Hendon 1990; Dunkerton and Baldwin 1995). Serra et al. (2008) applied a spectral filtering to OLR data at 10°N, 95°W to capture westward-propagating synoptic-scale disturbances with periods of several days. They showed that southwest–northeast-oriented waves exist, and the convection is located north and west of the cyclonic centers in the southerlies. Roundy and Frank (2004) showed an interesting snapshot around 160°–120°W, with a large-scale disturbance over the equator and a pair of positive and negative anomalies detected with total precipitable water (TPW) along the axis of the ITCZ. Other studies also showed that vortices are generated over the ITCZ associated with the barotropic instability (Nieto Ferreira and Schubert 1997; Wang and Magnusdottir 2005, 2006; Wang et al. 2010). Magnusdottir and Wang (2008) showed composites with zonally elongated westward-propagating relative vorticity at 850 hPa over the EP, using wavenumber–frequency-filtered 850-hPa vorticity over 8°–11°N, 130°–110°W as an index.

Utilizing First Global Atmospheric Research Program (GARP) Global Experiment (FGGE) Level III-b data, Tai and Ogura (1987) reported that disturbances over the EP are associated with divergence in the lower–middle troposphere and in the upper troposphere, corresponding to bipeaks of upward velocity observed around 850 and 250 hPa. On the other hand, the disturbances over the WP are observed to have divergence only in the upper troposphere, corresponding to a peak of upward velocity around 400–300 hPa (Reed and Recker 1971). The reason why the vertical structures are different in these two regions has not yet been fully elucidated.

In recent years, the dominance of shallow rain and congestus over the EP has been emphasized in some studies (Nesbitt et al. 2000; Berg et al. 2002; Kubar et al. 2007; Back and Bretherton 2009b) and in a companion paper (Yokoyama and Takayabu 2012, hereafter Part I). It also has been shown that a shallow meridional return flow exists over the EP in addition to a deep meridional return flow (Trenberth et al. 2000; Zhang et al. 2004, 2008; Nolan et al. 2007, 2010). One of the significant differences between the EP and the WP is the depth of the large-scale convergence field. Over the EP, there is a shallow (1000–925 hPa) convergence field accompanied by lower-tropospheric divergence just above it. This shallow convergence field is driven by strong sea surface temperature (SST) gradients (Back and Bretherton 2009a). In contrast, a relatively deep (1000–400 hPa) convergence field appears over the WP. Part I found that both shallow rain from congestus with rain top heights (RTHs) below 8 km and moderately deep rain from organized rain systems with RTHs in the range of 8–14 km are dominant over the shallow convergence region over the EP during the boreal autumn, while deep rain from small systems and very deep rain (RTHs ≥ 14 km) from organized systems are dominant in the WP warm pool region. It is possible that such differences in characteristics of rain systems are associated with a difference in the synoptic-scale disturbances.

Especially, it remains as a question how moderately deep rain is maintained over the EP. As discussed by Part I, the EP may be less favorable for the development of deep convection than the WP, because of 1°–2°C lower SST and divergence in the lower to middle troposphere. In this study, we will discuss a mechanism how organized rain systems are maintained in relation to the synoptic-scale disturbances over the EP.

In addition, it is worth analyzing the behavior of shallow rain from congestus in relation to the synoptic-scale disturbances statistically and quantitatively. Khouider and Majda (2006) developed a model that considers the roles of congestus, stratiform, and deep convective cumulus clouds in the dynamics of the convectively coupled equatorial waves. They emphasized that congestus play a role as a precondition for the development of deep clouds by moistening the lower troposphere. OLR and IR brightness temperature, which are often utilized to capture deep convection coupled with disturbances in many studies, may not be appropriate to represent variations in congestus. In this study, we also utilize TPW retrieved from satellite data to examine synoptic-scale disturbances associated with shallower disturbances.

We will also focus on differences between the EP and the WP in terms of energetics. Over the WP, latent heat released by deep convection contributes to the generation of the kinetic energy of disturbances (e.g., Nitta 1970, 1972; Reed and Recker 1971; Lau and Lau 1992). Similarly, Serra et al. (2008) showed that disturbances are primarily driven by latent heat released by deep convection around 6°–15°N, 95°W over the EP. Serra et al. (2010) also found that barotropic conversions contribute to the easterly wave energetics over the far eastern Pacific. However, the kinetic energy of the synoptic-scale disturbances propagating westward across the region around 6°–10°N, 140°–120°W, where much shallow rain is observed as well as much moderately deep rain in the boreal autumn (Part I), has not yet been well known. Energetics in this region may be different from those over the WP and over the far eastern Pacific.

The objectives of this study are to describe characteristics of synoptic-scale disturbances over the EP, and to examine the relationship among mesoscale rain systems, synoptic-scale disturbances, and the large-scale environment over the EP. There is a possibility that congestus also plays an important role in determining the characteristics of synoptic-scale disturbances over the EP. In this study, we analyze the synoptic-scale disturbances in the boreal autumn, when both congestus and organized rain systems are dominant over the EP ITCZ.

2. Data

We utilize National Oceanic and Atmospheric Administration (NOAA) interpolated OLR data (Liebmann and Smith 1996), and Special Sensor Microwave Imager (SSM/I) TPW data over the ocean (Wentz 1997; Wentz and Spencer 1998). Wind speeds and directions at a height of 10 m over the ocean surface are obtained from the Quick Scatterometer (QuikSCAT) data (Wentz et al. 2001). The 25-yr Japanese Re-Analysis (JRA-25; Onogi et al. 2007) data and Japan Meteorological Agency (JMA) Climate Data Assimilation System (JCDAS) data are also used. We additionally utilize the European Centre for Medium-Range Weather Forecasts (ECMWF) Interim Re-Analysis (ERA-Interim; Dee et al. 2011) data for comparison with results from JRA-25/JCDAS data.

We also use the Tropical Rainfall Measuring Mission (TRMM) latent heating research product, which is based on the spectral latent heating (SLH) algorithm (Shige et al. 2004, 2007, 2008). The algorithm uses TRMM Precipitation Radar (PR; 13.8 GHz) 2A25 product to retrieve heating profiles utilizing lookup tables for three rain types: convective, shallow stratiform, and deep stratiform rain. These lookup tables are derived from tropical precipitation, and latent heating data simulated by a cloud-resolving model. The lookup tables are based on rain top heights for convective rain and shallow stratiform rain, while it is based on rain strengths at the melting level for deep stratiform rain. In this study, Q1QR data are utilized from the product, where Q1 is the apparent heat source (Yanai et al. 1973) and QR is the radiative heating. We accumulated orbital Q1QR data and the total number of observational pixels at every 1.25° grid on a daily time scale.

The analysis period is boreal autumn from 2000 to 2005. In section 5 and part of section 4, the analysis period is extended to boreal autumn from 1998–2007, in order to increase the sampling. In this study, we especially focus on the difference in synoptic-scale disturbances between the EP and the WP. Hereafter, we refer to the longitudes of ~150°–100°W and ~130°–170°E as the EP and the WP, respectively. These regions are corresponding to those defined by Part I.

3. Methodology

The fast Fourier transform (FFT) is utilized to examine the abundance of synoptic disturbances statistically and quantitatively. We first obtain the anomalies by removing the mean and the trend in boreal autumn for each year, then applying a cosine tapering to the first and last 10% of the anomaly time series. Then, FFT is applied to the time series of anomalies to calculate power spectra for each year. After the simple moving mean, which uses the adjacent five data points centered on each point, the power spectra are averaged for 6 yr from 2000 to 2005. In addition, a bandpass filter (BPF) with half-power frequency cutoffs at 2.5 and 10 days is utilized to capture the disturbances, because spectral peaks are found at a period of 3–7 days (see section 4).

A composite analysis is utilized to elucidate the disturbance structures. BPF-vertical pressure velocity at 850 hPa (ω850) at 7.5°N, 130°W is used as the reference variable over the EP. The 83 reference days (DAY0s) are defined as the dates when the peak BPF ω850 exceeds one standard deviation (−0.092 Pa s−1) during boreal autumn 2000–05. This same threshold value is also used for composites of boreal autumn 1998–2007. In addition, we use BPF-vertical pressure velocity at 300 hPa (ω300) at 7.5°N, 150°E as the reference data over the WP, where ω300 varies more pronouncedly than ω850.

In section 4c, an energy budget analysis is utilized to examine the budget of the eddy kinetic energy. The energy budget equations are as follows:
e1
e2
e3
e4
e5
where K′ is the eddy kinetic energy and R is the gas constant for dry air. For an arbitrary variable X, denotes the 3-month average, and X′ denotes the anomaly obtained by subtracting and the linear trend from X. In addition, denotes the average taken over a period of the wave. The first (advKe), second (KmKe), third (GKe), and forth (AeKe) terms on right-hand side in Eq. (1) represent the advection of K′ by the environmental flow, the barotropic conversion from the mean kinetic energy to K′, the divergence of eddy geopotential flux, and the conversion from the eddy available potential energy to K′, respectively. The fifth term (D) represents the dissipation or subgrid-scale effects.

4. Characteristics of synoptic-scale disturbances over the eastern Pacific

a. Spectral characteristics

Figure 1 shows Hovmöller diagrams of OLR, TPW, and JRA-25/JCDAS ω850 and ω300 over the tropical Pacific for a period from 1 September to 30 November 2000. These variables are averaged over different latitudinal bands to capture their large variations at a period of 3–7 days (as will be shown in Fig. 3). For all variables, synoptic-scale westward propagations with a period of several days are found.

Fig. 1.
Fig. 1.

Hovmöller diagrams of (a) OLR (W m−2), (b) TPW (mm), (c) JRA-25 ω300 (Pa s−1), and (d) JRA-25 ω850 (Pa s−1) during 1 Sep–30 Nov 2000. OLR and ω300 are averaged between 5° and 12.5°N, TPW are averaged between 5° and 7.5°N, and ω850 are averaged between 5° and 10°N. Black crosses indicate the reference days (DAY0s) in the composite analysis over the EP. In (a), contours indicate MRG wave–filtered OLR anomalies, which are extracted in a similar way to Wheeler and Kiladis (1999), except that an antisymmetric–symmetric decomposition is not used. Contours are drawn at −10, −5, 5, and 10 W m−2, and negative values are contoured in dotted lines. The white slanted lines in (b) are missing data.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

Westward-propagating OLR signals over the WP generally have lower values, and are broader, compared to those over the EP (Fig. 1a). The pattern of ω300 is similar to that of OLR (Fig. 1c), which suggests that the variation in ω300 is related to the disturbances associated with deep convection. Over the WP, values of OLR and ω300 are lower during October and at the end of November, when convection is active associated with the Madden–Julian oscillation (MJO; Madden and Julian 1972). Note that any MJO-filtered data are not shown here. During early-to-mid-September, lower values of OLR and ω300 are observed over the EP when convection associated with the MJO is present. Some tropical cyclones also contribute to these large areas of less positive OLR and negative ω300 over the EP and the WP. Examining Hovmöller diagrams of OLR and ω300 filtered in the band of 2.5–10 days, it is found that the westward propagations of these signals are more pronounced over the WP around 150°E than over the EP around 130°W (not shown).

On the other hand, westward propagations of TPW signals at the period of several days are found over the entire tropical Pacific, and even over the EP, where the westward propagations of OLR are more obscure than the WP (Fig. 1b). Westward propagations of ω850 signals are also dominant over the EP (Fig. 1d). The pattern of ω850 is in good agreement with that of TPW, with the clearest disturbances around 130°W for both cases. Interestingly, there is an eastward progression of ω850 at the phase speed of ~2 m s−1 from 180°–160°W on 1 September to 110°–90°W on 16 October. This progression is on the same time scale as a downwelling oceanic Kelvin wave, when compared to the dataset constructed by Roundy and Kiladis (2007), which tends to couple to convection over the Pacific (Gribble-Verhagen and Roundy 2010). Further studies are needed to reveal whether there are interactions between ω850 and oceanic Kelvin waves.

Figure 2 shows power spectra of the above four variables, which are computed at 7.5°N, 150°E for the WP and at 7.5°N, 130°W for the EP. Over the WP, a spectral peak of OLR (~400 W2 m−4) is detected at a period of ~4–5 days. A spectral peak of JRA-25/JCDAS ω300 (~8 × 10−3 Pa2 s−2) is also found at a similar periodicity (Fig. 2e, black line). In contrast to OLR and ω300, TPW and ω850 have smaller spectral power at all periods with no spectral peaks.

Fig. 2.
Fig. 2.

Power spectra of (a),(b) OLR (W2 m−4); (c),(d) TPW (mm2); (e),(f) ω300 (10−3 Pa2 s−2); and (g),(h) ω850 (10−3 Pa2 s−2) in the boreal autumn during 2000–05. (a),(c),(e),(g) The WP (7.5°N, 150°E) and (b),(d),(f),(h) the EP (7.5°N, 130°W). In (e)–(h), black and gray lines indicate JRA-25/JCDAS and ERA-Interim data, respectively. Abscissa indicates periods (day) in the logarithmic scale with 91 days on the extreme left and 0.5 day on the extreme right. Ordinate indicates power spectra multiplied by frequency. Error bars indicate the confidence intervals at the 95% significance level.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

On the other hand, over the EP, the spectral peaks of TPW (~9 mm2) and ω850 (~6 × 10−3 Pa2 s−2) are pronouncedly detected at a period of ~5 days. The power spectral distributions of TPW and ω850 are very similar to each other, which suggests that TPW varies with the convergence and divergence associated with the disturbances in the lower troposphere. Over the EP, the OLR also has a spectral peak of ~300 W2 m−4, although slightly smaller than the peak over the WP, at a period of ~5 days. In addition, ω300 has spectral power comparable to ω850 at a period of several days (Fig. 2f, black line).

We also compare power spectra for JRA-25/JCDAS data with those for ERA-Interim data. The power spectrum for ERA-Interim shows no significant peak in ω300 over the WP (Fig. 2e, gray line), in contrast to the power spectrum based on JRA-25/JCDAS (Fig. 2e, black line). Similarly, over the EP, the spectral power of ω300 for JRA-25/JCDAS is about 3 times larger than that for ERA-Interim (Fig. 2f). Since the spectral peak of OLR is detected over both regions, it is considered reasonable that ω300 has a spectral peak at the same periodicity with OLR, because vertical velocity in the upper troposphere should vary with deep convection. Note that ERA-Interim ω300 also captures variations in OLR associated with MRG wave–type disturbances over the EP (not shown). Hereafter, we utilize JRA-25/JCDAS data for the analyses because of a good agreement between ω300 and OLR. The discrepancies come from many differences in the process to produce the reanalysis datasets. The models used in JRA-25/JCDAS and ERA-Interim have a spectral T106 horizontal resolution and 40 vertical layers and a TL255 resolution and 60 layers, respectively. Different physical parameterization, especially for cumulus convection, may also be important for the discrepancies. The discrepancies are currently under investigation. On the other hand, the power spectra of ω850 obtained from the two reanalysis data are similar to each other.

Next, we integrate the power spectra of the above four variables from 3 to 7 days to examine their horizontal distribution (Fig. 3). The spectral power of OLR is large over both the WP and the EP. Large spectral power expands to the west from ~130°W, corresponding to the passages of synoptic-scale disturbances. To the west of ~160°E, the power extends west-northwestward, which is consistent with the west-northwestward propagation of the TD-type disturbances over the WP (Takayabu and Nitta 1993). The power spectra of ω300 are in good agreement with those of OLR.

Fig. 3.
Fig. 3.

Spectral power of (a) OLR (W2 m−4), (b) TPW (mm2), (c) ω300 (10−3 Pa2 s−2), and (d) ω850 (10−3 Pa2 s−2) integrated from the period of 3.03–7.00 days in the boreal autumn during 2000–05.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

On the other hand, the spectral power of TPW and ω850 are dominant over the EP. For TPW, there is a minimum of power at the center latitudes of the ITCZ around 7°–12.5°N, where OLR and ω300 have the largest variation. Instead, there are two bands of large power on both the southern and northern edges of the ITCZ. This result is consistent with Roundy and Frank (2004), who find latitudinally dual centers of disturbances in TPW during boreal autumn. The power of ω850 is larger around the southern band of TPW than the northern band.

These results indicate that synoptic-scale disturbances over the EP have significant variations in vertical velocity both in the upper and lower troposphere. It suggests that these disturbances are associated not only with deep convection but also with shallow convection. On the other hand, over the WP, the disturbances are primarily associated with deep convection and variations in vertical velocity in the upper troposphere, but less with the lower-tropospheric variations represented by TPW and ω850.

b. Composite structure

A composite analysis is performed to examine structures of synoptic-scale disturbances. The 83 cases are used for the EP during boreal autumn 2000–05. The DAY0s for the EP in 2000 are plotted with cross marks in Fig. 1. Note that DAY0s are not necessarily in good agreement with MRG wave–filtered OLR anomalies (Fig. 1a, contours), which are extracted in a similar way to Wheeler and Kiladis (1999), except that an antisymmetric–symmetric decomposition is not used. In addition, strongly negative ω850 during the end of September are not sampled, because they are largely contributed by variations at periods longer than 2.5–10 days.

First, a composite of BPF-horizontal winds at 850 hPa is shown in Fig. 4a. Interestingly, the disturbance has both a cyclonic vortex with its center around ~9°N and a large-scale cross-equatorial circulation. Hereafter, we refer to the disturbances with this structure as “coupled disturbances.” The cyclonic vortex is accompanied by a weaker anticyclonic vortex to the east. In addition, cross-equatorial southerlies at the longitudes from 135° to 115°W and weaker northerlies to the west are found. This coupled structure is clearly shown for a few days at least (not shown). The coupled disturbance is confirmed by comparing the JRA-25/JCDAS wind composite at 1000 hPa with that of ocean surface wind anomalies obtained from QuikSCAT observational data (Figs. 4b,c). Compared to those at 850 hPa, the cross-equatorial winds are less significant both with 1000-hPa JRA-25/JCDAS winds and with the QuikSCAT data. It is suggested that while the vortex structure is significant to the near surface, the cross-equatorial winds are more significant in the lower troposphere than in the boundary layer.

Fig. 4.
Fig. 4.

Composites of JRA-25/JCDAS BPF-winds (m s−1) at (a) 850 and (b) 1000 hPa over the EP at DAY0 in the boreal autumn during 2000–05. (c) As in (a), but for QuikSCAT wind anomalies (m s−1) at 10 m over the ocean. The composites are based on all 83 cases. Arrows represent composite winds where either the zonal or meridional component is significant at the 95% level.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

There may be a concern that the coupled disturbance could artificially result from combining dates with a vortex and dates with a cross-equatorial circulation. Therefore, we investigated whether individual cases of such coupled disturbances really exist over the EP. Consequently, some snapshots of a coupled disturbance at 850 hPa were confirmed. Cases only with a vortex and cases only with a cross-equatorial circulation were also found. We consider that a vortex and a cross-equatorial circulation can sometimes coexist.

To elucidate what effects the coexistence produces, we subjectively selected 21 coupled disturbance cases among all 83 cases. The cases with both a vortex and a cross-equatorial circulation at 850 hPa are selected in the region of 20°S–20°N, 180°–80°W. In most cases, the vortex is located around 5°–15°N, 135°–115°W. Tropical cyclones are not included in the vicinity of the reference point. Cross-equatorial southerlies tend to be found to the south or southeast of the vortex, and cross-equatorial northerlies are to the west of the southerlies. The composite for the 21 cases clearly show the coupled disturbance, which has more pronounced cross-equatorial northerlies compared to those shown in Fig. 4a.

Second, composite longitude–pressure cross sections of BPF-meridional winds at 7.5°N and at the equator are shown in Fig. 5. Note that composite figures shown in the remaining part of this section are for all cases, unless otherwise stated. At 7.5°N, both southerlies and northerlies have their maxima at ~925 hPa, and slightly tilt toward the east with altitudes. Significant northerlies are found from the surface to ~400 hPa, while significant southerlies are only found under ~700 hPa. At the equator, both southerlies and northerlies, the latter of which are strongest at ~500 hPa, have larger tilts toward the east with altitudes. Northerlies are piled up on southerlies. The vertical structure of meridional winds at the equator is similar to that of an MRG wave–type disturbance over the central Pacific shown by Takayabu and Nitta (1993).

Fig. 5.
Fig. 5.

Composite longitude–pressure cross sections of BPF-meridional winds (m s−1) at 850 hPa over the EP at DAY0 in the boreal autumn during 2000–05. Cross sections at (a) 7.5°N and (b) the equator are shown. The composite is performed for all 83 cases. Shades indicate significant regions at the 95% level.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

Third, Fig. 6 indicates composite Hovmöller diagrams of BPF-meridional velocity at 850 hPa obtained from JRA-25/JCDAS data. The zonal wavelengths are estimated at ~4000 km and ~8000 km at 7.5°N and at the equator, respectively. These values are consistent with the horizontal structure shown in Fig. 4. Meridional wind signals propagate westward with a phase velocity of ~9 m s−1 at 7.5°N. At the equator, meridional wind signals propagate westward with a phase velocity of ~20 m s−1, which is much faster than that at 7.5°N. These values are also consistently estimated with composites for meridional ocean winds obtained from the QuikSCAT (Figs. 6b,d).

Fig. 6.
Fig. 6.

Composite Hovmöller diagrams of (a),(c) JRA-25/JCDAS BPF-meridional winds (m s−1) at 850 hPa and (b),(d) QuikSCAT meridional wind anomalies (m s−1) at 10 m over the ocean in the boreal autumn during 2000–05. Cross sections are at (a),(b) 7.5°N and (c),(d) the equator. The composite is performed for all 83 cases. In (a) and (c), the contour interval is 0.2 m s−1, and values greater than or equal to 0 m s−1 are contoured in solid lines. In (b), contours are set at −1, −0.5, 0.5, and 1 m s−1. In (d), contours are drawn at −0.4, −0.2, 0.2, and 0.4 m s−1. Shades indicate significant regions at the 95% level. Thick dashed lines in (a) and (c) indicate the phase speeds of −9 and −20 m s−1, respectively.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

Moreover, Fig. 7 shows a composite Hovmöller diagram of meridional winds at the equator, which is composed only of the 21 cases with distinct coupled disturbances. The characteristics of meridional winds are more obvious in this figure than Fig. 6c. The estimated wavelength and the phase velocity at the equator are ~8000 km and ~−20 m s−1, respectively. From 110°W to the date line or more westward, there is a westward-propagating signal of meridional winds. In addition, wave packets propagate eastward with a group velocity of ~5 m s−1, which means that the energy of the cross-equatorial circulation originates to the west of 150°E, and propagates eastward.

Fig. 7.
Fig. 7.

As in Fig. 6c, but for a composite from DAY-7 to DAY7 for the 21 cases with the coupled disturbance of a vortex and a cross-equatorial circulation. The contour interval is 0.4 m s−1 and values greater than or equal to 0 m s−1 are contoured in solid lines. Thick solid and dashed lines indicate the phase speed of 20 m s−1 and the group velocity of 5 m s−1, respectively.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

It is worth noting that the phase velocity of the disturbance at 7.5°N differs from that at the equator. It suggests that a vortex and a cross-equatorial circulation are basically independent from each other. However, the two disturbances may propagate together as a coupled disturbance, once they encounter each other. They may enhance each other through a coupling of convection, resulting in the coexistence of them.

These characteristics of the cross-equatorial circulation are similar to those of an MRG wave–type disturbance. Liebmann and Hendon (1990) showed the MRG wave–type disturbance with a zonal wavelength of ~6700 km and a phase velocity of −18 m s−1 over the eastern-central Pacific. Takayabu and Nitta (1993) also observed the dominance of MRG wave–type disturbances with a zonal wavelength of ~8000 km near the date line. In addition, the cross-equatorial circulation in this study has the similar vertical structure of meridional winds at the equator, as already shown in Fig. 5b.

Here, we examine to what extent we can understand the cross-equatorial circulation as an MRG wave. Its intrinsic frequency calculated from the composite results is plotted against zonal wavenumbers over the dispersion relation curves (no shown). To consider the Doppler effect, we utilize the intrinsic frequency, which is calculated by subtracting the 3-month mean zonal velocity (~−2.4 m s−1) at 850 hPa at the reference point. There is a significant correlation between the phase velocity of the disturbance and wind velocity at 850 hPa (Liebmann and Hendon 1990). We also consider the effect of cumulus convection on the phase velocity of the disturbance, which is coupled with convection, by using the value at 7.5°N. Because it is considered that convectively coupled equatorial waves are excited by cumulus convection, the place and the structure of the disturbance at the next step are determined by where convection occurs next. On the other hand, the disturbances strongly affect the place where convection occurs. Therefore, it may be appropriate to consider advection by the mean zonal wind at 7.5°N, where the disturbance is coupled with convection.

As a result, the intrinsic frequency of ~0.2 day−1 is calculated against a wavenumber of 5. The value is plotted near the dispersion relation curve of an MRG wave with an equivalent depth of 30 m. This result is similar to the Takayabu and Nitta (1993)’s result for the disturbances from 160°E to 180° in the boreal summer. From the dispersion relation of an MRG wave with an equivalent depth of 30 m, the group velocity is estimated as ~8 m s−1. After subtracting the mean zonal wind velocity from it, the group velocity, which can directly be compared with the group velocity obtained from Fig. 7, is estimated as 6 m s−1. Thus, the group velocity of 5 m s−1 estimated from Fig. 7 is comparable to the group velocity of an MRG wave with an equivalent depth of ~30 m. Therefore, the cross-equatorial circulation is consistently identified as an MRG wave–type disturbance.

Next, composite longitude–pressure cross sections of BPF-divergence and BPF-vertical pressure velocity at 7.5°N for the EP and for the WP are compared (Fig. 8). These composites are performed for all cases. Over the WP, the composite disturbance has deep convergence in the layer of 1000–300 hPa and divergence in the layer of 300–100 hPa. The upward velocity peak is found at ~300 hPa. Over the EP, there are two sets of convergence and divergence: convergence in the layers of 1000–800 and 500–300 hPa, and divergence in the layers of 800–500 and 300–100 hPa. Correspondingly, there are two upward velocity peaks at ~800 and ~300 hPa over the EP. Hereafter, the vertical structure shown over the EP is referred to as the double-deck structure.

Fig. 8.
Fig. 8.

Composite longitude–pressure cross sections of (a),(b) BPF-divergence (s−1) and (c),(d) BPF-vertical pressure velocities (Pa s−1) at DAY0 in the boreal autumn during 2000–05. (a),(c) For the EP and (b),(d) for the WP. The composite is performed for all cases. In (a) and (b), contours are set at −5, −3, −2, −1, 0, 1, 2, 3, 5, 10, and 15 × 10−6 s−1, and values greater than or equal to 0 s−1 are contoured in solid lines. In (c) and (d), contours are set at −0.14, −0.1, −0.06, −0.02, 0, and 0.02 Pa s−1, and values greater than or equal to 0 Pa s−1 are contoured in solid lines. Shades indicate significant regions at the 95% level.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

Figure 9 shows composite time series of TRMM Q1QR profiles near the reference point. The composite is performed in the boreal autumn from 1998 to 2007. Over the EP, both deep heating with its peak at ~7.5 km and shallow heating with its peak at ~2 km coherently vary with time with the shallow peak slightly ahead of the deep peak. In other words, it is observed that both deep and shallow heating are modulated by a disturbance. Small variations in shallow heating over the EP result from the westward propagation of the disturbance in the ITCZ with abundant congestus. Note that our analysis period consists mostly of the La Niña and normal phases of El Niño–Southern Oscillation, and the shallow heating tends to be less associated in case of the El Niño phase. In contrast, over the WP, the separation of the lower and the upper heating as seen over the EP is not clearly observed.

Fig. 9.
Fig. 9.

Composite time series of TRMM Q1QR profiles (a) over the EP (6.25°–8.75°N, 131.25°–128.75°W) and (b) over the WP (6.25°–8.75°N, 148.75°–151.25°E) in the boreal autumn during 1998–2007. The composite is performed for all cases. Contours are set at 1, 1.3, 1.6, 1.9, 2.2, 2.5, 3, 5, and 7 K day−1.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

In addition, the heating profile for the WP is antisymmetric, while that for the EP is symmetric. It is suggested that the disturbance over the WP is associated with transition from shallow convection to sporadic deep convection to organized convective systems. Over the EP, the disturbance is basically associated with shallow convection, and deep convection tends to occur only around DAY0 when the environment is favorable for deep convection, as will be shown in section 5.

c. Energy budget

Figure 10 shows composite vertical profiles of energy budget terms averaged over the EP (5°–15°N, 140°–120°W) and the WP (5°–15°N, 140°–160°E). In both regions, AeKe makes the largest contribution to the energy generation of the disturbance at 250 hPa. The largest AeKe over the EP (~40 × 10−5 m2 s−2) is smaller than that over the WP (~65 × 10−5 m2 s−2). The converted energy is redistributed up and down by GKe. Over the EP, the second peak of AeKe (~15 × 10−5 m2 s−2) is found at 850 hPa. The converted energy at 850 hPa is also redistributed up and down in the same manner as in the upper troposphere.

Fig. 10.
Fig. 10.

Composite vertical profiles of each term (10−5 m2 s−3) on the right-hand side of the budget equation of the eddy kinetic energy in Eq. (1). Profiles are averaged over (a) the EP (5°–15°N, 140°–120°W) and (b) the WP (5°–15°N, 140°–160°E). Black solid, black dashed, gray solid, and gray dashed lines indicate the conversion from the kinetic energy of the 3-month-averaged field to the eddy kinetic energy (KmKe), the conversion from the eddy available potential energy to the eddy kinetic energy (AeKe), the convergence of the geopotential fluxes (GKe), and the advection of the eddy energy by the mean flow (advKe), respectively. The composite is performed for all cases.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

Moreover, the eddy available potential energy generated through diabatic heating (GAe) is calculated. At 850 and 250 hPa, values of GAe are similar to those of AeKe (not shown). Therefore, AeKe is the conversion from the eddy available potential energy, which is mostly generated by cumulus convection, to K′. In other words, over the EP, both deep convection and congestus generate kinetic energy of the disturbance in the upper and in the lower troposphere, respectively.

In addition, there is a peak of KmKe (~7 × 10−5 m2 s−2) at 925 hPa over the EP, which is in contrast to the WP, where KmKe is ~5 × 10−5 m2 s−2 or less through the layer of 1000–400 hPa. After examining the breakdown for KmKe at this level, the first and second terms on the right-hand side in Eq. (3) mostly contribute to KmKe (not shown). The first and second terms represent the shear conversion term and the energy amplification term through the meridional convergence, respectively. The energy amplification through the meridional convergence is larger than the shear conversion, and corresponds to the spectral power of ω850 from 3 to 7 days (Fig. 3d). Thus, the shallow boundary layer convergence field over the EP amplifies the eddy kinetic energy at 925 hPa through the barotropic conversion. This term is the largest around 9°N at 925 hPa, which agrees with the position where the vortex of the coupled disturbance has the largest meridional winds. Thus, it is suggested that this barotropic conversion term generates the energy for the vortex of the coupled disturbance.

5. Rain characteristics in relation to the synoptic-scale disturbances

In this section, we analyze TRMM-based Q1QR and divergence profiles to examine the relationship between cumulus convection and the coupled and double-deck structure of the disturbance over the EP. We use 35 cases with distinct coupled disturbances during boreal autumn from 1998–2007 for the composites.

Figure 11a shows composite BPF-ω850 and BPF-ω300 of the coupled disturbance with contours. Vertical pressure velocity at 850 hPa is widely distributed throughout the vortex with its maximum around 6°–8°N, 128°–132°W. On the other hand, ω300, which is smaller than ω850, has its maximum to the northeast of the maximum ω850. The ratio of BPF-ω850 to BPF-ω300 shows that there is a clear contrast between the southwestern and the northeastern parts of the vortex, which are divided at around the longitude of 130°W (Fig. 11a, shades). Updraft at 300 hPa is more dominant in the northeastern part of the vortex than in the southwestern part.

Fig. 11.
Fig. 11.

(a) Composite BPF-ω850 (black contours) and BPF-ω300 (blue contours) of the coupled disturbance over the EP at DAY0 in the boreal autumn during 1998–2007. The contours start from −0.04 Pa s−1 with an interval of −0.04 Pa s−1. Shades indicate the ratios of composite BPF-ω850 to BPF-ω300. Shades are plotted where both BPF-ω300 and BPF-ω850 are significant with the 95% level. Arrows indicate horizontal winds (m s−1) at 850 hPa where either the zonal or meridional component is significant at the 95% level. (b) Profiles of TRMM Q1QR (K day−1). Gray and black lines are for the profiles averaged over the southwestern part (6.25°–7.5°N, 135°–130°W) and the northeastern part (7.5°–8.75°N, 131.25°–125°W, and 7.5°–10°N, 123.75°W, and 8.75°–10°N, 122.5°–121.25°W) of the vortex, respectively. These analysis regions are indicated with gray boxes in (a). Error bars indicate confidence intervals with the 95% significance level. (c) As in (b), but for vertical pressure velocity profiles. The composite is performed for 35 cases with the coupled disturbance.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

Figures 11b,c show Q1QR profiles and vertical pressure velocity profiles, respectively. Profiles are averaged in the southwestern and in the northeastern part of the vortex (the western and the eastern box in Fig. 11a, respectively). These two parts are defined based on the ratio of 1.6 in Fig. 11a. Note that Q1QR and vertical pressure velocity are not bandpass filtered in Figs. 11b,c. For both the southwestern and the northeastern part of the vortex, Q1QR profiles have two peaks of heating at heights of ~2.5 and ~7.5 km. The shallower peak is ~2.5 K day−1 in the southwestern and the northeastern part, and there is no significant difference between the two parts of the vortex. In contrast, the deeper peak is significantly greater in the northeastern part than in the southwestern part. These results are consistently and even more pronouncedly shown in the vertical pressure velocity profiles. Interestingly, the northeastern part of the vortex is the region where cross-equatorial southerlies of the MRG wave–type disturbance flow into the vortex (Fig. 11a). It is suggested that the variation of rain characteristics is closely related to the coupled structure of a vortex and an MRG wave–type disturbance.

Next, a longitude–pressure section of BPF-divergence profiles is examined in relation to the coupled disturbance over the EP. Figure 12 shows divergence profiles averaged over 6.25°–10°N. The composite is performed for 35 cases with coupled disturbances. Similar to the structure shown in Fig. 8a, the coupled disturbance has the double-deck structure with divergence in the upper and the middle to lower troposphere. Divergence at 150 hPa, which is found around 125°W, is located about 5°E of divergence at 700 hPa. This result is consistent with the dominance of deep heating in the northeastern part of the vortex.

Fig. 12.
Fig. 12.

As in Fig. 8a, but for a composite of 35 cases with the coupled disturbance over the EP in the boreal autumn during 1998–2007. The contour interval is 1 × 10−6 s−1 and values greater than or equal to 0 s−1 are contoured in solid lines. The cross section is averaged between 6.25° and 10°N.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

Figures 13a–d show composite latitude–pressure cross sections of BPF-meridional circulation and BPF-divergence averaged over every 5° of longitude. In the region of 140°–135°W, weak shallow convergence exists with very weak divergence above the convergence. Both the lower-tropospheric convergence and the middle–lower-tropospheric divergence increase in the region of 135°–130°W. In the same region, significant convergence and divergence exist aloft. The upper-tropospheric divergence further increases in the regions of 130°–125°W and 125°–120°W, which is consistent with the result that deep heating is large in the northeastern part of the vortex. The upper-tropospheric divergence is centered at higher altitudes in the regions of 125°–120°W and 130°–125°W than 135°–130°W. At the same time, the lower-tropospheric convergence and the middle–lower-tropospheric divergence are weaker in the region of 125°–120°W.

Fig. 13.
Fig. 13.

Composite latitude–pressure cross sections of BPF-divergence (shades; 10−6 s−1) over the EP at DAY0 in the boreal autumn during 1998–2007. Cross sections are averaged over (a) 140°–135°W, (b) 135°–130°W, (c) 130°–125°W, and (3) 125°–120°W. Purple contours indicate significant regions at the 95% level. Arrows indicate meridional and vertical wind velocity, and black arrows indicate vectors where either the zonal or meridional component is significant at the 95% level. Vertical velocities are stretched to appear clearly in the figures. The composite is performed for 35 cases with the coupled disturbance.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

Most notably, Fig. 12 shows that the convergence in the eastern part of the vortex (1000–700 hPa) around 120°W is deeper than the convergence in the western part (1000–925 hPa) around 140°W. Cross-equatorial southerlies in the lower troposphere are also deeper in the northeastern part of the vortex than in the southwestern part (Fig. 13). As already shown in Fig. 5, the cross-equatorial southerlies of the MRG wave–type disturbance tilt eastward with altitudes, which is consistent with the deeper convergence in the northeastern part of the vortex. These results suggest that deep heating in the northeastern part of the vortex is the result of cross-equatorial southerlies; producing deep convergence in the northeastern part of the vortex.

Finally, we compare composite BPF-divergence profiles of the isolated vortex (9 cases) and the isolated MRG wave–type disturbance (13 cases) with those of the coupled disturbance (Fig. 14). These cases are subjectively selected. The isolated vortex has more significant divergence in the layer of 700–600 hPa than in the upper troposphere. On the other hand, the isolated MRG wave–type disturbance has significant divergence only around 200 hPa. Thus, the vortex and the MRG wave–type disturbance are associated with the middle to lower-tropospheric divergence and the upper-tropospheric divergence, respectively. As a result of the superposition of the two disturbances, the coupled disturbance may have divergence peaks both in the middle–lower and in the upper troposphere. According to the positions of the two disturbances, the divergence profiles of the coupled disturbance vary with longitudes. In addition, the isolated MRG wave–type disturbance has a smaller upper-tropospheric divergence region than the coupled disturbance, and has a slightly lower maximum in upper-tropospheric divergence. It is suggested that the coexistence of the vortex and the MRG wave–type disturbance as a coupled disturbance may serve as a mutual intensification mechanism.

Fig. 14.
Fig. 14.

As in Fig. 12, but for composites of the cases with (a) the isolated vortex-type disturbances and the cases with (b) the isolated MRG wave–type disturbances.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

6. Summary and discussion

a. Characteristics of synoptic-scale disturbances

In this study, we examined characteristics of synoptic-scale disturbances over the EP during boreal autumn, and compared them with those over the WP. First, over the EP, spectral peaks of TPW and ω850 are detected at the periods of 3–7 days. Spectral peaks in the same periodicity range are also detected with OLR and ω300, although these powers are smaller than those over the WP. These results indicate that westward-propagating synoptic-scale disturbances over the EP are associated with congestus as well as with deep convection. On the other hand, over the WP, variations in OLR and ω300 are dominant. Contrastingly, the spectral peaks of the variables in the lower troposphere such as TPW and ω850 are not found in the WP.

Second, composite analysis revealed that the coupled disturbances, which have both a southwest–northeast-oriented vortex with a center at ~9°N and an MRG wave–type disturbance, exist over the EP. The vortex has a zonal wavelength of ~4000 km, propagating westward at a phase speed of ~9 m s−1, and slightly tilts eastward with altitudes. The vortex has the wavenumber of 10 and the intrinsic frequency of ~0.2 day−1, which are similar to those for so-called easterly waves. However, it has shallower characteristics compared to previous studies. The vortex around 130°W tends to have the lower–middle-tropospheric divergence. In addition, the vortex is the largest meridional winds at ~925 hPa, which is lower compared to previous studies. Reed and Recker (1971) shows that the altitude with the largest meridional winds of easterly waves gradually decreases from the WP to the central Pacific, which suggests that it is even lower around 130°W. The altitude around 130°W is lower than that around 112.5°W (Tai and Ogura 1987).

On the other hand, the MRG wave–type disturbance has a zonal wavelength of ~8000 km, propagating westward with a phase speed of ~20 m s−1, and largely tilts eastward with altitudes. A wave packet associated with the MRG wave–type disturbance propagates eastward at the group velocity of ~5 m s−1.

Since the phase speed of the vortex differs from that of the MRG wave–type disturbance, these two disturbances are considered as basically independent from each other. Actually, cases with only a vortex and cases with only an MRG wave–type disturbance are detected. It is suggested that a vortex and an MRG wave–type disturbance, which basically exist independently of each other, can coexist as a coupled disturbance under proper conditions. The energy of the MRG wave–type disturbance, which propagates eastward from the west of the reference point, can excite a new MRG wave–type disturbance around the vortex. Once the vortex is coupled with the MRG wave–type disturbance, they can enhance each other while moving together for a considerable time (a few days at least).

Vertically, the composite disturbance over the EP has a double-deck structure, which has pairs of convergence and divergence both in the upper and in the middle-lower troposphere, associated with upward velocity maxima both in the upper and in the lower troposphere. The composite disturbance over the WP has a single pair of deep convergence and divergence, and has a single upward velocity maximum in the upper troposphere.

Furthermore, both over the EP and over the WP, the conversion from the eddy available potential energy to the eddy kinetic energy (AeKe) at 250 hPa shows the largest contribution to the energy generation of the disturbances. AeKe at 850 hPa is additionally found over the EP. Thus, over the EP, both deep convection and congestus generate kinetic energy of disturbances in the upper and in the lower troposphere, respectively. Furthermore, disturbances over the EP have deep and shallow peaks of observationally based heating, which are pronouncedly separated from each other.

b. The environment favorable for the coupled disturbances

Here, we consider why the coupled disturbances are dominant over the central-eastern Pacific. Over the EP, the shallow boundary layer convergence field amplifies the eddy kinetic energy at 925 hPa, where the meridional winds of the vortex are largest, through the barotropic conversion (KmKe). Therefore, it is considered that the vortex is dominant in the strong shallow convergence field over the EP. On the other hand, as suggested by Hendon and Liebmann (1991), the symmetric distribution of high SSTs along the equator may be favorable for the MRG wave–type disturbances. SST is relatively symmetric against the equator to the west of ~140°W during boreal autumn, while a low SST region extends in the Southern Hemisphere to the east of ~140°W. Consistently, the values of MRG-filtered OLR variance increase westward from the EP with the maximum over the central Pacific in the Northern Hemisphere (Fig. 19g of Roundy and Frank 2004). The similar relationship between SST and MRG-filtered OLR variance is also found during boreal spring (not shown). Thus, the central-eastern Pacific is considered favorable both for the vortex and for the MRG wave–type disturbances.

c. The relationship among the synoptic-scale disturbances, cumulus convection, and the large-scale environment

The shallow heating with a peak at the height of ~2.5 km has no significant difference between the southwestern and the northeastern part of the vortex of the coupled disturbance. On the other hand, the deep heating with a peak at ~7.5 km is significantly greater in the northeastern part of the vortex. It is found that the coupled MRG wave–type disturbance is providing a cross-equatorial southerly flow to the northeastern part of the vortex. The depths of the southerly winds entering the vortex and their convergence are deeper in the northeastern part of the vortex. As described in section 1, the EP is less favorable for the development of deep convection because of relatively low SST and the shallow convergence field. Our results suggest that the deep heating is maintained with the existence of the deep convergence associated with the coupled-structure disturbances in the shallow convergence field over the EP.

It is also interesting to discuss the coupled disturbance in terms of the shallow meridional circulation (e.g., Zhang et al. 2008; Nolan et al. 2010). The meridional circulations of the coupled disturbance (Fig. 13) are consistent with the depths of cumulus convection associated with the disturbance. The ratio of shallow heating to deep heating is larger in the southwestern part of the vortex than the northeastern part, which is corresponding to stronger northerly shallow (700–400 hPa) return flows in the southwestern part. Deep convection is dominant with the deep convergence in the northeastern part, where the northerly shallow return flows are weak.

Figure 15 summarizes the relationship among cumulus convection, the synoptic-scale coupled disturbances, and the large-scale environment over the EP in the boreal autumn. Over the EP, the shallow convergence enhances shallow rain from congestus. Westward-propagating vortex-type disturbances, which modulate the shallow convergence, are found over the EP. The vortex deepens through receiving the energy from congestus and the shallow convergence field. Around the vortex, the eastward-propagating energy of the MRG wave–type disturbance excites a new MRG wave–type disturbance. On the other hand, relatively deep cross-equatorial southerly winds associated with the MRG wave disturbance flow into the northeastern part of the vortex and produce deep convergence. Thus, rain systems are well organized through the deep convergence produced by the coupled disturbances, which results in the dominance of moderately deep rain including convective rain and much stratiform rain at the northeastern part of the vortex. Deep rain plays a role of the main energy source for the coupled disturbances.

Fig. 15.
Fig. 15.

Schematic for the relationships among cumulus convection, the synoptic-scale coupled disturbances, and the large-scale environment over the EP in the boreal autumn. Small and large circles indicate a vortex disturbance and an MRG wave–type disturbance, respectively. The thick light gray arrow represents deep convergence associated with the MRG wave–type disturbance. Dark gray arrows indicate shallow convergence, which is largely driven by the strong SST gradient. Black solid and open arrows indicate phase and energy propagation, respectively. Small and large cloud-shaped figures are depicted as congestus and well organized systems, respectively.

Citation: Monthly Weather Review 140, 9; 10.1175/MWR-D-11-00251.1

On the other hand, the WP in the boreal autumn is favorable for deep rain because of high SST and the deep convergence field. There, deep rain can be directly coupled with the synoptic-scale disturbances with a deep structure, without the production of deep convergence by the coupled disturbances. Besides, small deep rain systems such as cumulonimbi are often caused by unstable stratification over the high SST, and can exist without synoptic-scale disturbances. After all, the organization of rain systems with the help of the coupled disturbances is important to bring deep rain to the EP, where the environment is less favorable for developing deep convection, compared to the WP.

Acknowledgments

This work is based on a part of the first author’s doctoral dissertation. The first author thanks Drs. M. Kimoto, M. Satoh, H. Nakamura, K. Iga, H. Hasumi, and A. Sumi at the University of Tokyo for their helpful comments and discussions. This work is supported by Precipitation Measuring Mission project of the Japan Aerospace Exploration Agency (JAXA). The authors would like to express their gratitude to three anonymous reviewers for their very helpful comments. They also want to acknowledge the Japan Meteorological Agency, the European Centre for Medium-Range Weather Forecasts, and NOAA for providing valuable data. SSM/I data and QuikSCAT data were produced by Remote Sensing Systems, and are available online (http://www.remss.com).

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