1. Introduction
Warm rain refers to the rainfall from shallow clouds that do not participate in the ice-phase processes (Byers and Hall 1955; Battan and Braham 1956; Austin et al. 1996; Petty 1999; Johnson et al.1999). Warm rain is formed primarily by the coalescence of water droplets of different sizes as they fall at different terminal velocities within the clouds (Baker 1993; Schumacher and Houze 2003) and is prevalent in marine clouds (Berg et al. 2002). As an integral component of the tropical precipitation systems, shallow marine clouds contribute an appreciable amount to the total precipitation in the tropics (Short and Nakamura 2000; Schumacher and Houze 2003; Lau and Wu 2003; Liu and Zipser 2009). In addition, since warm rain often encompass large areas, it is a significant fraction of total rain areas (Austin et al. 1996). Thus, warm rain plays an important role in the heating and moistening of the lower troposphere, as well as the planetary radiation budget (Johnson and Lin 1997; Johnson et al. 1999; Medeiros et al. 2008).
Warm rain systems usually come in the form of small isolated convection (Schumacher and Houze 2003). However, some studies found that shallow clouds can also occur as cloud streets, clusters, or mesoscale arcs (Malkus and Riehl 1964; Warner et al. 1979; Nair et al. 1998). Recent studies based on the Rain in Cumulus over the Ocean (RICO) experiment documented that local rain rates exceeding 1 mm h−1 are not infrequent in the trades and are often associated with arc-shaped cloud clusters (Rauber et al. 2007a,b; Minor et al. 2011; Zuidema et al. 2012). By using snapshots of precipitation systems from the Tropical Rainfall Measuring Mission (TRMM; Kummerow et al. 1998) satellite, Liu and Zipser (2009, 2013) found that mesoscale warm rain systems with strong radar echoes do occur in some coastal regions and tropical oceans. Figure 1 shows a case of a warm precipitating system observed over the eastern Pacific ITCZ, with a size of about 1180 km2 and maximum near-surface reflectivity of 36 dBZ but a low radar echo top (3.7 km) below the freezing level. The minimum infrared brightness temperature over this system is 274 K. Systems like this are not uncommon over the tropical eastern Pacific.
An example of TRMM-observed warm organized RPFs with warm cloud tops and a large precipitation area in the eastern Pacific on 21 January 2002. (a) The TRMM PR near-surface reflectivities (color fill) and VIRS TB11 (gray shades). (b) The PR echo-top heights (color fill) and daily sea surface temperature (gray shades and contours). Note that this RPF, at the center of the panels, has a minimum infrared brightness temperature TB11 value (274 K) > 0°C and PR echo-top height ~3.3 km below freezing level, and an area of ~1180 km2; and is located between warm and cold SST regions.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
Although large-size shallow clouds are observed, the mesoscale organization of shallow precipitating systems received limited attention. Malkus (1957) argued the importance of sea surface temperature (SST) anomalies as a factor in the clustering of shallow cumulus clouds. Another important conceptual idea to explain clustering is preconditioning—that is, clouds modifying their local condition in such a way that the conditions are favorable for the subsequent formation of the secondary clouds (Randall and Huffman 1980; Khairoutdinov and Randall 2002; Kuang and Bretherton 2006). Observations from RICO and large-eddy simulations have provided evidence that the formation of cold pools plays a role in triggering a new shallow convection in the trade wind regime (Zuidema et al. 2012; Seifert and Heus 2013).
The different characteristics of convections are closely related to large-scale environments (Del Genio and Kovan 2002). Some relevant studies have emphasized the roles of various feedbacks in the organization of deep convective systems. They include the positive feedbacks between water vapor and convection (Held et al. 1993; Tompkins 2001), surface flux and radiation (Emanuel 1986; Nilsson and Emanuel 1999; Raymond 2000; Bretherton et al. 2005; Stephens et al. 2008; Muller and Held 2012), wind and surface heat exchange (e.g., Neelin et al. 1987; Emanuel 1987; Brown and Bretherton 1995; Chao and Deng 1996), and gravity wave adjustment (e.g., Mapes 1993; Oouchi 1999) and cold-pool dynamics (Tompkins 2001; Jeevanjee and Romps 2013). To improve the convective parameterization in climate models, it is imperative to understand the complicated interactions between convection and large-scale environment.
This paper attempts to use a large number of satellite-observed samples to investigate the relationship between shallow organized precipitating systems and their associated large-scale environments. The two main objectives of this paper are 1) to examine the geographical distribution of shallow organized rain systems, their seasonal, diurnal variations, and various properties (sections 3a and 3b) and 2) to investigate the typical favorable environments in order to promote horizontal organization of shallow convective systems (sections 3c and 3d). These environmental conditions include specific thermodynamic variables, large-scale atmospheric circulation, and sea surface temperature (SST). The satellite observations, atmospheric mean state reanalysis datasets, and methodology are described in section 2. Results are discussed in section 3 and a summary is given in section 4.
2. Data and methods
a. TRMM precipitation feature datasets
In this study, 16-yr radar precipitation feature (RPF) data (Liu et al. 2008) based on the TRMM version 7 products are utilized to examine the characteristics of precipitation systems. RPFs are identified by grouping pixels using the TRMM 2A25 measured nonzero near-surface rain rate. TRMM 2A25 is a product containing the orbital data of radar reflectivity and precipitation rate profiles observed by the TRMM Precipitation Radar (PR; Iguchi et al. 2000, 2009). Variables describing the properties of RPF have been derived (e.g., geocenter location, number of rain pixels, and rain volume). The size of an RPF is calculated as the total pixel count multiplied by the size of each pixel within 17.96 km2 (before the satellite orbit boost in August 2001) or 20.35 km2 (after boost). The coldest cloud-top temperature is inferred from the minimum infrared brightness temperature at 10.8-μm wavelength (TB11) by the TRMM Visible and Infrared Scanner (VIRS). We use rain area and cloud-top temperature to define the warm organized systems. Note that the TRMM PR has a minimum sensitivity limitation of 18 dBZ, which causes it to miss a large portion of weak precipitation in the tropics and underestimate warm rainfall (Lebsock and L’Ecuyer 2011; Liu and Zipser 2014). In this study, we focus on the warm rain systems that are detectable by TRMM PR, which excludes most of the light precipitation from stratocumulus clouds. However, to help assessment of weak rainfall that is missed by PR, the precipitation features defined by grouping raining areas detected by the TRMM Microwave Imager (TMI; Kummerow et al. 2001, 2011) are also used. More details on TRMM precipitation feature datasets are introduced by Liu et al. (2008).
The warm RPFs in the tropics are selected by specifying the RPFs with minimum infrared brightness temperature warmer than 273 K. More than 27 million warm RPFs are defined over 20°S–20°N from 16 yr of TRMM data (Table 1); 25 million are located over tropical oceans and 2.2 million are over tropical land. The selected RPFs are “pure” warm rain systems, which eliminates the ones attached or embedded in cold clouds as discussed by Liu and Zipser (2009). Most of the defined warm RPFs are isolated with a small horizontal extent. Figure 2 shows the rainfall fractions from RPFs over tropical oceans and tropical land, binned by rain area and minimum TB11. About 78% of warm rain (RPFs with minimum TB11 > 273 K) over tropical oceans is contributed by RPFs with an area smaller than 200 km2 (vertical dashed line), and such warm precipitation features are defined as warm shallow isolated systems hereafter. However, there is still a small fraction (2.3%) of warm rain that is contributed by RPFs with a rain area greater than 500 km2. These are then classified as shallow organized RPFs. These defined warm organized RPFs are barely observed over tropical land. It is worth mentioning that the value of 500 km2 is arbitrarily chosen to distinguish these isolated warm rain systems. We have repeated the analysis for warm systems greater than 200 and 1000 km2; conclusions on the general properties of large-scale environments for these warm organized RPFs are basically the same.
Mean properties of warm RPFs (minimum TB11 > 273 K) and warm organized RPFs (minimum TB11 > 273 K and size > 500 km2) in tropical (20°S–20°N) land and ocean.
Volumetric rainfall fraction over (a) tropical oceans and (b) tropical land areas from RPFs binned by rain area and minimum infrared brightness temperature at 10.8-µm wavelength (TB11). The values within each panel add up to 100%. The ranges to define shallow isolated (SI), shallow organized (SO), and deep organized (DO) RPFs are shown by the horizontal dashed lines in (a) and the vertical dashed lines correspond to rain areas of 200, 500, and 2000 km2.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
b. Large-scale environmental variables
To examine the characteristics of large-scale mean states associated with warm organized RPFs, atmospheric fields from the European Centre for Medium-Range Weather Forecasts (ECMWF) interim reanalysis (ERA-Interim; Dee et al. 2011) are adopted in this study, including relative humidity (RH), temperature (T), horizontal wind (u and υ components), and pressure vertical wind (ω). These variables are available on a 1.5° × 1.5° horizontal grid at 37 different pressure levels with 6-h intervals. Horizontal moisture flux divergence (often referred to simply as moisture divergence by the forecasting community) is calculated from specific humidity and horizontal velocity. The NOAA high-resolution (0.25° × 0.25°) blended analysis of daily sea surface temperature (Reynolds et al. 2007) is used to describe oceanic surface conditions. To derive the environment conditions for each RPF, these variables are first temporally interpolated to the time of RPF. Then, values of variables at the center location of RPFs are selected from the nearest grid. Because there are too many warm RPFs, we only derived the environmental variables for RPFs with at least 4 PR pixels, about 80 km2 in size.
3. Results
Table 1 lists the sample mean properties of warm RPFs and warm organized RPFs over tropical oceans and land between 20°S and 20°N from 16 yr (1998–2013) of TRMM data. The population of warm RPFs is 10 times larger over tropical oceans than over tropical land. More than 0.1 million warm organized RPFs exist over tropical oceans, nearly 25 times the population of those over tropical land. The rainfall contribution from warm RPFs is larger over tropical oceans than over tropical land. Warm RPFs contribute around 9.2% of the total rainfall over ocean but only 1.8% over land. Compared to warm isolated RPFs, warm organized RPFs have larger mean rain areas (~770 km2), higher surface rain rates, and greater maximum near-surface reflectivity (~37 dBZ). The maximum surface rain rate from warm organized RPFs may reach about 11–12 mm h−1, heavier than that from small-size warm RPFs (~2.5 mm h−1). The mean maximum echo-top height of warm organized RPFs is around 4 km, lower than the typical freezing level of about 4.5 km in tropics.
a. Distribution, variation, and shape of warm organized systems
Figure 3a shows the geographical distributions of unconditional monthly rainfall from warm RPFs of all sizes between 20°N and 20°S, averaged over 1998–2013. Unconditional rainfall means the total rainfall averaged over all times regardless of whether it is raining, as opposed to conditional rain rate, which is given by averaging only at times of precipitation. In general, the rainfall amount from warm RPFs is quite small over tropical land (~1.8% of total rain in Table 1). Most of the warm rainfall occurs over tropical oceans, including the eastern Pacific intertropical convergence zone (ITCZ), the eastern part of the South Pacific convergence zone (SPCZ), the south Indian Ocean, the central North Pacific, Hawaii, and several coastal regions, such as the eastern coasts of Madagascar, eastern Brazil, Costa Rica, the Philippines, and the Caribbean Islands. Warm RPFs generate less than 20 mm month−1 rainfall over most of these regions. The geographical distribution of warm rainfall is consistent with Fig. 3a in Liu and Zipser (2009). Figure 3b shows the total unconditional rainfall (color fill) against rainfall (contours) from warm organized RPFs with cloud tops warmer than 273 K and areas larger than 500 km2. The regions with a large amount of rainfall roughly correspond to deep mesoscale convective systems (MCSs). Rainfall from warm organized RPFs occurs in the eastern Pacific ITCZ, the eastern part of the SPCZ, the south Indian Ocean, the central North Pacific, Hawaii, and several coastal regions, which are boxed in Fig. 3c. The sample mean properties of warm organized RPFs in these regions are listed in Table 2. The maximum near-surface rain rate in warm organized systems is about 10–15 mm h−1. Rain rates are higher in coastal regions of the Northern Hemisphere (e.g., Costa Rica and the Caribbean) than in the eastern Pacific. Also notable in Fig. 3b is that with the exception of the eastern Pacific ITCZ, most warm organized RPFs occur over regions with large-scale subsidence and few deep heavy rain systems. The warm organized precipitating system over the eastern Pacific ITCZ is found to be at the southern edge of the narrow ITCZ. For example, the warm organized RPF over the eastern Pacific shown in Fig. 1 is centered between warm and cold SST regions. Figure 3c shows the percent of warm rainfall contributed by the warm organized RPFs. Warm organized RPFs over coastal regions and large-scale subsidence regions like the SPCZ and south Indian Ocean contribute about 12%–14% (Table 2) of the local warm rain, a relatively greater percentage than that over the eastern Pacific ITCZ (9%).
(a) Geographical distribution of monthly mean rainfall (color fill, mm month−1) from warm RPFs. (b) Geographical distribution of mean unconditional rainfall (mm month−1) from TRMM PR 3A25 product (color fill) and from organized warm RPFs (contours). (c) Percent of warm rainfall from warm organized RPFs (color fill). These distributions are created using 2° × 2° boxes.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
Mean properties of warm organized RPFs over selected regions.
The PR had a minimum detectable signal of 18 dBZ before the satellite orbit boost in 2001. According to the TRMM Z–R relationship for convective rainfall near the surface, 18 dBZ corresponds to a minimum rain rate of 0.1–0.5 mm h−1. Thus, the PR may miss drizzle and light rain composed of low concentrations of small raindrops that produce weak reflectivity signatures or shallow enough echo-top heights that are obscured by the radar surface return (Short and Nakamura 2000; Berg et al. 2006, 2010; Lebsock and L’Ecuyer 2011). Berg et al. (2010) used collocated TRMM PR and CloudSat observations to estimate that the PR misses 10% of the total rainfall accumulation in the tropical and subtropical oceans. To address such light warm rain, TMI-detected precipitation features (TPFs) are used. TMI precipitation retrieval utilizes the low-frequency channels (10–19 GHz) that are sensitive to the liquid water in the atmosphere; it is relatively more sensitive to weak precipitation (such as drizzle in the stratocumulus) than PR (Liu and Zipser 2014). Warm organized TPFs are defined by TPFs with minimum infrared brightness temperature warmer than 273 K and area of the TPFs greater than 500 km2. Roughly, the regions with a large amount of monthly mean rainfall from warm organized TPFs shown in Fig. 4a correspond well with those from RPFs in Fig. 3b. However, more unconditional rainfall from warm organized systems is observed and the associated raining regions are 19% larger. The large mean rainfall (Fig. 4b) from warm organized TPFs missed by PR is prevalent in marine stratocumulus regions, including the eastern tropical North Pacific, South Pacific, and South Atlantic. These regions with climatological subsidence feature persistent low clouds with drizzle and light rainfall. Figure 4c indicates that PR could miss 30%–60% of rainfall from warm organized convective systems detected by passive microwave retrievals in these regions. However, considering that rainfall quantity derived from RPFs is more reliable than that from TPFs, in this study we focus on the warm precipitation systems that are detectable by PR. To address the properties of the cases missed by PR, we have conducted a similar large-scale environment analysis with TPFs. The results are similar to those using RPFs and thus are not shown.
(a) Geographical distribution of monthly mean unconditional rainfall (mm month−1) from TRMM PR 3A12 product (color fill) and rainfall from warm organized TPFs (contours). (b) Rainfall from warm organized TPFs but missed by PR. (c) Percent of rainfall from warm organized TPFs missed by PR.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
The seasonal variations of unconditional rainfall from warm organized RPFs are shown in Fig. 5. A larger amount of warm organized rainfall occurs in winter than in summer. Most of these wintertime warm rains are probably related to the enhanced inversion by the advection of the stable air by the trade winds (Schubert et al. 1995). Some extratropical influences may be involved as well, especially over the southern SPCZ (Kiladis and Weickmann 1992).
(top to bottom) Seasonal variation of unconditional rain rate (color fill, mm month−1) from warm organized RPFs with minimum VIRS TB11 > 273 K and size > 500 km2. The distributions are created using 2°× 2° boxes. The sample biases have been removed by using the number of pixels sampled in each box.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
According to the regions and possible factors leading to these organized warm systems, we may roughly divide them into three groups. The first group occurs over the ocean near the coasts. The shallow organized RPFs in this group are likely associated with the inversion caused by the advection of stable conditions by the trade winds and the forcing by topography similar to the line convection near the windward coast of Hawaii (Austin et al. 1996; Carbone et al. 1998; Wang and Chen 1998). The second group occurs over deep oceans, including the eastern SPCZ, the central-south Indian Ocean, and the central North Pacific. There are large-scale subsidences over these trade wind regions; warm organized rain systems are associated with the sea surface temperature and the strength of the subsidence. The third group occurs in the tropical ITCZ, dominantly over the central and eastern Pacific. It should be interesting to study the different mechanisms of these systems over different regions. Since warm organized RPFs dominantly occur over the eastern Pacific in the tropical ITCZ, we will focus on warm organized systems over the eastern Pacific ITCZ at 0°–10°N and 160°–100°W.
b. Tropical eastern Pacific
To compare warm isolated precipitation systems, warm organized precipitation systems, and deep cold MCSs over the eastern Pacific ITCZ, warm isolated RPFs are selected with cloud tops warmer than 273 K and areas within 80–200 km2. The cold organized RPFs are defined with the minimum infrared brightness temperature colder than 210 K and rain area larger than 2000 km2. The characteristics of warm and cold organized systems over the eastern Pacific ITCZ are listed in Table 3. Cold organized RPFs are roughly 3 times more populated than warm organized RPFs. The former contributes 62% of the total rainfall in the tropical eastern Pacific, while the latter contributes only 0.64% of the total rainfall. Cold organized RPFs generate more intense rainfall, larger mean raining area, larger maximum near-surface reflectivity, and colder minimum infrared brightness temperature than warm organized RPFs. The average maximum echo-top height for cold organized RPFs is about 12 km.
Mean properties of warm organized (TB11min > 273 K and size > 500 km2) and cold organized (TB11min < 210 K and size > 2000 km2) RPFs in tropical the eastern Pacific (0°–10°N, 160°–100°W).
Figure 6a shows the seasonal variation of the population of warm organized RPFs (red solid line), warm RPFs (red dashed line), and all RPFs (black solid line) over the eastern Pacific ITCZ. The population of warm organized RPFs shows a larger seasonal variation than warm isolated and all other RPFs. The greater number of warm organized RPFs occur in boreal winter (DJF) rather than in boreal summer (JJA) (Fig. 6a). The seasonal variations of the rainfall generated by warm organized RPFs (red solid line), warm RPFs (red dashed line), and all RPFs (black solid line) are displayed in Fig. 6b. Warm rain over the eastern Pacific and the rainfall from warm organized RPFs show a similar seasonal variation, with a larger amount of rain that occurs in boreal winter when there is a stronger shallow meridional circulation (Zhang 2004; Yokoyama et al. 2014). In contrast, total rainfall undergoes a different annual cycle, and more precipitation occurs between May and August, with the peak in May (Fig. 6b). The diurnal variations of the population of warm RPFs, warm organized RPFs, and all RPFs over the eastern Pacific are shown in Fig. 6c. Since the diurnal cycle of rainfall over the ocean is almost completely due to an increase of the number of systems, instead of rain intensity (Nesbitt and Zipser 2003), the cycle of rainfall (not shown) is similar to that of their counts. The diurnal cycle of population for warm RPFs or all RPFs is weak. All RPFs exhibit similar diurnal cycles, with peaks at 0300 LT and minima at 1000 LT, consistent with the result in Nesbitt and Zipser (2003). With limited samples, it is difficult to determine the most frequent local time for warm organized RPFs. They do seem to peak later in the morning than the warm RPFs shown in Fig. 6c. However, this is inconclusive without evidence from additional samples.
(a) Monthly variation of population of warm organized RPFs (solid red), warm RPFs (dashed red), and all RPFs (solid black) over the eastern Pacific ITCZ (0°–10°N, 160°–100°W). Fraction is the percent of the population in each month divided by total population in all months. (b) As in (a), but for the variation of unconditional rainfall from the three groups of RPFs. (c) As in (a), but for the diurnal variation of the population.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
Following the same method used by Nesbitt et al. (2006) and Liu and Zipser (2013), the shapes of RPFs are described by using the ratio of minor and major axes of ellipses fitted to raining regions. A smaller ratio implies a more linear shape of RPF. Figure 7 shows the histogram of ratios for warm organized and cold organized RPFs over the eastern Pacific. Compared to cold organized systems, the ratio of minor/major axes for warm organized RPFs is smaller than the ratio for cold ones. The peak value is 0.4 for cold organized RPFs but is only 0.2 for warm organized RPFs. This indicates that warm organized RPFs are more likely to have a linear shape. Table 3 shows that 34% of warm organized RPFs have ratios smaller than 0.3, in comparison with only 17% of cold organized RPFs. The warm organized RPFs over other regions are also more linear, as listed in Table 2. Over most of these regions, 30%–40% of warm organized RPFs have a minor/major ratio less than 0.3. Vietnam has the highest fraction of linear-shaped warm organized systems among all selected regions. The linear-shaped warm precipitation was also observed during the RICO experiment (Rauber et al. 2007a,b). Usually, warm rain systems are isolated because the cold pool from the short-lived precipitation could stabilize the surroundings and prevent new convections from occurring at the same location. The cold pool would more likely lead to a propagating system with a linear shape at the edge. In addition, a linear-shaped system is sustained relatively more easily if it is due to moisture feeding from one side with low-level wind shear, such as squall lines. However, all these factors do not explain why warm organized systems tend to have more-linear shapes than cold systems. Further study on this is warranted.
The histogram of the linearity of RPFs for cold organized RPFs (blue) and warm organized RPFs (red) over the eastern Pacific ITCZ. The linearity of features is calculated using the ratio between minor and major axes of an ellipse fitting the feature region.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
c. Large-scale environments of warm organized RPFs over the eastern Pacific ITCZ
The various characteristics of precipitation systems could be closely related with the different large-scale environmental conditions. The mean atmospheric states are first examined in the vertical structure of thermodynamic properties for warm isolated and warm organized RPFs.
1) Vertical profiles of thermodynamic variables
Composite profiles of RH (%), pressure vertical velocity ω (Pa s−1), moisture divergence (kg kg−1 s−1), and horizontal wind convergence (s−1) for warm isolated RPFs (black) and warm organized RPFs (red) over the eastern Pacific ITCZ are plotted in Fig. 8. Filled regions denote the standard deviations for each corresponding variable. Generally, the RH profile has a surface value of 85% over the eastern Pacific and decreases to 20% at 400 hPa (Fig. 8a). Then, a sharp increase in relative humidity occurs above 400 hPa. The profile is consistent with climatology of relative humidity profiles over tropical oceans (Liu et al. 1991). However, the magnitudes of RH for warm isolated RPFs and warm organized RPFs are different. The significant difference happens in the mid- to upper troposphere from 700 up to 200 hPa such that warm organized RPFs accompany a drier mean state. The midtropospheric dry structure has also been found for deep convective systems to impact the evolution of convection (Mapes and Zuidema 1996; Roca et al. 2005; Zuidema et al. 2006; Takemi 2006). The RH at the lower troposphere is slightly higher for warm organized RPFs, but the difference is far less pronounced compared to that at the mid- to upper level (Fig. 8a). The ω profile in Fig. 8c shows the low-level large-scale ascent with a peak at 850 hPa and the weak descent above 600 hPa for warm organized RPFs. The dryness at the midtroposphere for warm organized RPFs and the corresponding weak subsidence indicates that the subsiding free troposphere is too dry and warm to maintain buoyancy for deep convection even if other favorable conditions are satisfied. Consistently, the lower troposphere displays a stronger shallow (1000–800 hPa) convergence for warm organized RPFs than that for warm isolated RPFs (Fig. 8d). In addition, surface flux is considered to be an environmental factor affecting the evolution of convection (Bretherton et al. 2005). The profiles in Fig. 8b show the enhanced near-surface moisture convergence for warm organized RPFs.
(a) Composite relative humidity (%) at times and locations of warm isolated (black) and warm organized RPFs (red) using ERA-Interim. (b) Moisture divergence (kg kg−1 s−1), (c) vertical pressure velocity (Pa s−1), and (d) horizontal wind divergence (s−1). Color shaded areas denote the standard deviation for each variable. The statistical estimation of the mean value passes the significance test 99% level. To emphasize the values at low levels, the pressure levels are shown in a linear instead of a traditional log scale.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
The mean environmental conditions for warm isolated, warm organized, and cold organized RPFs in the eastern Pacific ITCZ are listed in Table 4. Large and cold (deep) systems correspond to strongest near-surface moisture convergence and highest RH at both low level (at 850–1000 hPa) and midlevel than warm systems. Cold systems tend to occur at warmer SSTs (27.8°C) than warm isolated and organized systems. Note that warm organized RPFs are associated with the largest SST gradient. Additionally, compared to warm isolated RPFs, warm organized systems over the eastern Pacific tend to occur in a moister planetary boundary layer, drier midtroposphere, stronger mid- to upper-level subsidence, and a more enhanced moisture convergence near the surface. The averaged large-scale environment conditions for warm isolated and warm organized RPFs in other regions are also tested (not shown). The similar vertical structures are generally observed for warm organized RPFs over different regions, except that the warm organized RPFs over the tropical eastern Pacific have a stronger low-level ascent and moisture convergence and relatively weaker midlevel subsidence than other regions.
Large-scale environment conditions for warm isolated, warm organized, and cold organized RPFs in the eastern Pacific (0°–10°N, 160°–100°W).
Environmental factors do not impact convection solely but rather jointly with other factors. To consider the variation of precipitating systems corresponding to a combination of two environmental variables, 2D histograms (contour) are created for warm RPFs, categorized by near-surface moisture divergence versus vertical pressure velocity at 850 hPa ω850 (Fig. 9a) and relative humidity against vertical velocity at 500 hPa ω500 (Fig. 9c). The fractions (color fill) of warm RPFs having an area larger than 500 km2 are calculated at each bin, which indicates the probability for warm RPFs to become organized. Similarly, 2D histograms are shown for cold organized RPFs in Figs. 9b,d. The majority of RPFs are associated with a relativelyweak near-surface moisture convergence and a weak updraft at low level (contours in Figs. 9a,b). Precipitating systems more likely become organized when there is a stronger low-level large-scale ascent and more enhanced moisture convergence (color fills in Figs. 9a,b). Figure 9c shows that warm organized RPFs are more likely to occur when there is low midlevel moisture, regardless of whether the midtroposphere is ascending or descending (color changes more horizontally than vertically). This confirms the role of midtroposphere dryness in encouraging shallow convective systems into larger sizes. Note that a warm system can occur under both positive and negative ω500 when midlevel RH is low, indicating that the dry air is not always caused by the large-scale descent. The source of dry air could be from horizontal advection (Zuidema et al. 2006). In contrast to warm organized RPFs, cold RPFs are more likely organized when midlevel RH gets moister (Fig. 9d).
(a) A 2D histogram of warm RPFs with minimum VIRS TB11 > 273 K (contours) over the eastern Pacific categorized by near-surface moisture divergence (kg kg−1 s−1) and 850-hPa vertical pressure velocity ω850. The fraction of the being warm organized RPFs with an area > 500 km2 in each bin is shown in color fill. Bin sizes are 0.5 × 10−7 kg kg−1 s−1 for moisture divergence and 0.05 Pa s−1 for ω850. (c) As in (a), but categorized by midlevel (700–300 hPa) RH (%) and 500-hPa vertical pressure velocity (ω500, Pa s−1). Bin sizes are 2% for RH and 0.025 Pa s−1 for ω500. (b),(d) As in (a),(c), but for cold organized RPFs with minimum VIRS TB11 < 210 K and area > 2000 km2; the total population of RPFs is displayed with contours and color fill denotes the population fraction of cold organized RPFs.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
2) Large-scale circulation
The composite meridional–pressure cross section of vertical pressure velocity and the meridional large-scale circulation corresponding to warm isolated (Fig. 10a), warm organized (Fig. 10b), and cold organized RPFs (Fig. 10c) over the eastern Pacific ITCZ are shown. The x axis indicates the latitude relative to the composite center. The large-scale ascent is found at the center of the RPFs (Fig. 10c). Cold RPFs have an ascent up to 200 hPa. The ascent for warm RPFs decreases from the surface to near zero roughly at 500 hPa (Figs. 10a,b). To the north and south of this convective region for warm RPFs, there exists an asymmetric large-scale subsidence. To the south, the subsidence region in the cold tongue over the central and eastern Pacific corresponds to a strong large-scale descent in the whole troposphere about 10°–15° south of the warm RPFs. The downdraft to the south has two peaks; one is at about 700 hPa and another is at 450 hPa, consistent with Yokoyama et al. (2014). To the north of the convection, the downdraft is weaker than that at the south and there is only one peak at about 800 hPa. The shallow meridional circulations from the center to the south and north of the warm RPFs are consistent with past literature (Yin and Albrecht 2000; Zhang et al. 2004). To the south, both upper-tropospheric return flows around 250–150 hPa and shallow return flows around 700–500 hPa coexist for warm and warm organized RPFs. On average, the shallow return flows are stronger than upper-level return flows. To the north, southerlies are found throughout the layer above 800 hPa and northerlies below it. The return flows above 800 hPa increase with altitude. This profile of meridional wind at the north of convection is different from that in Yokoyama et al. (2014), who concluded that there is no shallow south return flows for shallow convection. The inconsistent result may be due to different region and season selection.
Composite meridional pressure cross section of vertical pressure velocity (Pa s−1, contours) and meridional circulation (m s−1, arrows) for (a) warm isolated RPFs, (b) warm organized RPFs, and (c) cold organized RPFs over the eastern Pacific (0°–10°N, 160°–100°W). Meridional and vertical winds are from the ERA-Interim. The x axis indicates the latitude relative to the composite center. (d) The difference of vertical pressure velocity ω (Pa s−1) corresponding to warm organized RPFs and warm isolated RPFs. Solid line denotes the zero vertical velocity. The difference in meridional velocity (m s−1) between (e) warm organized and warm isolated RPFs and (f) warm organized and cold organized systems. Solid line denotes the zero meridional velocity associated with warm organized PFs.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
Although the general structure of atmospheric circulation for warm organized RPFs and warm isolated RPFs are similar, their magnitudes are different. The differences in vertical pressure velocity and meridional winds associated with warm organized RPFs and warm isolated RPFs are shown in Figs. 10d and 10e, respectively. The solid line denotes the zero vertical velocity. The shallow meridional circulation to the south of the RPFs is significantly enhanced for organized warm RPFs, indicated by a stronger shallow meridional circulation with low-level [~(800–500) hPa] northerly flow, together with near-surface southerlies, ascending motions in ITCZ, and stronger descending motions south of the reference center. The low-level northerly flow is even stronger compared to cold organized RPFs (Fig. 10f). The shallow meridional cell north of the warm organized RPFs appears to be enhanced in comparison to isolated ones, but the difference is much less pronounced than the south branch.
3) SST
Previous studies demonstrate that convective variability is dependent on SST (Zhang 1993; Khairoutdinov and Emanuel 2010; Wing and Emanuel 2012, 2014). Large-scale atmospheric low-level convergence can be determined by the absolute values of local SST (Neelin et al. 1987). Deep convective heating especially intensifies when SST approaches ~28°C (Gadgil et al. 1984; Ramanathan and Collins 1991; Zhang 1993). Some literature emphasizes the importance of large-scale SST gradients in determining regional-scale ascent in the marine atmospheric boundary layer (Lindzen and Nigam 1987; Back and Bretherton 2009) when accompanied by large-scale ascent aloft. The elevated SST and SST gradients locally modify the atmospheric boundary layer and reduce tropospheric stability, thereby triggering and/or amplifying precipitation in these regions (Minobe et al. 2008; Xie 2004). The composited SST distributions relative to the center of warm isolated, warm organized, and cold organized RPFs over the eastern Pacific are shown in Fig. 11. RPF centers are marked with plus signs. The x axis and y axes are the latitudes and longitudes relative to the center of RPFs. Horizontal wind convergence (s−1) at the level of 950 hPa is overlaid with contours. Three types of precipitating systems correspond to different SST patterns. Warm organized RPFs are located at the southern edge of warm SST; at the north side of their convective center, there shows the strongest SST gradient. Warm isolated RPFs correspond to relatively warmer SST than warm organized RPFs, but the SST gradient northward is weakest. Cold organized RPFs have the warmest SST and secondary strong SST gradient on the north side of their center location. The SST gradient is important for the aggregation of shallow precipitating systems.
Composite SST (color fill) for (a) warm isolated, (c) warm organized, and (e) cold organized RPFs over the tropical eastern Pacific. SST is from NOAA high-resolution daily product and its spatial resolution is 0.25° × 0.25°. Plus markers (blue +) denote the RPF center locations. The x axis and y axis indicate the latitudes and longitudes relative to the center of each reference grid. Horizontal wind convergence (s−1) at 950 hPa is shown with contours. (b),(d),(f) The composite SST meridional gradient associated with the three types of systems in (a),(c),(e), respectively. The solid line denotes the zero SST gradient.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
A natural question pertaining to the warm organized RPFs in the tropical eastern Pacific can be raised: why is it observed only at the south side of the ITCZ? To answer this question, differences in the thermodynamical structures between both sides of the RPFs are explored. The composite meridional–vertical cross section of potential temperature θ, equivalent potential temperature θe (Fig. 12, top) and temperature, relative humidity, atmospheric circulation (Fig. 12, bottom) relative to the center location of warm isolated, warm organized, and cold organized RPFs are shown in Fig. 12. As expected, centers of three types of RPFs all correspond to the local maxima of near-surface moist static energy (MSE) indicated by high θe at levels below 900 hPa. The value of θe is largest for cold organized RPFs, moderate for warm isolated RPFs, and smallest for warm organized RPFs. The smaller value of MSE for warm organized RPFs is due to the relatively colder low troposphere indicated by lower θ (contour in Fig. 12c). For warm isolated RPFs, their convective centers correspond to the local maximum of both MSE and θ, while it is not true for warm organized RPFs. The latter’s convective center corresponds to maximum MSE, but the local maximum θ is at the north side of center, indicating the important contribution of surface moisture to the local MSE at convective center. Additionally, it is to be noted that there is a slightly lower temperature at the surface of the cold organized RPFs’ center, probably owing to the impact of precipitation evaporation considered by the ECMWF model (Nuijens et al. 2009).
Meridional pressure cross section of composite equivalent potential temperature (shaded, θe in K) and potential temperature (contours, θ in K) for (a) warm isolated with warm cloud top (TB11 > 273 K) but small area (<200 km2), (c) warm organized with warm cloud top (TB11 > 273 K) but large area (>500 km2), and (e) cold organized RPFs with cold cloud top (TB11 < 210 K) and large area (>2000 km2) over the tropical eastern Pacific. The three plots have the same vector scales and color bars. (b),(d),(f) Composite RH (colors, %) and meridional circulation (arrows; ω in Pa s−1 and υ in m s−1) for the three types of systems in (a),(c),(e), respectively. They share the same scales and color bars. As in Fig. 10, the x axis indicates latitude relative to the composite center. Thick solid lines in (b),(d),(f) denote the zero large-scale vertical velocity.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
The differences of relative humidity profiles (color fill in Figs. 12b,d,f) confirm the important roles of moisture structure in development of convection. The low troposphere (below 800 hPa) is moistest in the convective region of cold organized RPFs and driest for warm isolated RPFs characterized by lowest relative humidity at their centers. Meanwhile, the midlevel troposphere between 700 and 300 hPa, is driest for warm organized RPFs and has corresponding large-scale descent. Stronger shallow meridional circulations at the south side of the ITCZ (Fig. 10e) bring a significant amount of low-level moisture to the area sustaining a large area of warm RPFs. At the same time, a midlevel dryness and its corresponding large-scale strong subsidence suppress the convective systems into a shallow state. Deep convective systems may be encouraged only if the midlevel becomes moist (Fig. 12f) and suppression from entrainment of dry air is no longer significant, consistent with Del Genio and Kovan (2002).
An asymmetric meridional structure of thermodynamics is found in the deep ITCZ over the eastern Pacific (Figs. 12e,f). Namely, in comparison to the northern side of the reference convective center, the southern side corresponds to a moister troposphere at lower levels (below 800 hPa), higher surface equivalent potential temperature, drier mid- to high troposphere, and stronger large-scale descending at high level. The low-level favorable convective conditions and mid- to high-level suppression are likely the prevalent conditions for shallow organized convection at the southern edge of ITCZ. In contrast, the drier surface and weaker shallow meridional circulation on the northern side of ITCZ do not provide enough moisture convergence for convective plumes to sustain a large system area. The environmental variables used in this study are mainly from model products after assimilating some remote sensing observations. It is encouraging that the reanalysis data capture the center of the precipitation systems over the tropical eastern Pacific. However, results shown here still need further validation with more observations.
d. SPCZ
As in Fig. 10, the composited large-scale circulations for the warm organized and warm isolated RPFs in the SPCZ are shown in Fig. 13. In contrast to the zonally oriented ITCZ in the eastern Pacific accompanying a north–south convergence at the center (Fig. 10a), the large-scale circulation in the SPCZ is diagonal along the northwest and southeast direction. Additionally, the low-level updraft at the convective center is much weaker over the SPCZ than those over the eastern Pacific. The SPCZ shows an asymmetric meridional circulation at the northern and southern sides of the convective center. To the northern side, there is a strong large-scale subsidence, which goes through the whole troposphere and peaks around 800 hPa. This large-scale descent is the key to the formation of trade cumulus over the region. The downdraft on the southern side is much weaker than the one at the northern side. The large-scale meridional circulation displays the same pattern for warm organized RPFs (Fig. 13a) and warm isolated RPFs (Fig. 13b), while the warm organized RPFs have a relatively stronger ascent at low levels and much stronger subsidence at mid- to high levels (Fig. 13c)
As in Fig. 11, the composited SST and meridional SST gradient for the warm organized and cold organized RPFs in the SPCZ are shown in Fig. 14. The warm organized RPFs in the SPCZ are located between the southeast extension of warm SST over the South Pacific and the large-scale subsidence region over the cold SST tongue. The corresponding SST for warm organized RPFs in the SPCZ is around 27.3°C, lower than the SST (~28.4°C) associated with cold organized RPFs. Similar to the result for tropical eastern Pacific, the center of warm organized RPFs over the SPCZ corresponds to a stronger SST gradient, in comparison to cold organized systems.
As in Fig. 11, but for (a),(c) composite SST and (b),(d) SST meridional gradient for warm organized and cold organized RPFs over the SPCZ. Solid line denotes the zero SST gradient.
Citation: Journal of Climate 29, 9; 10.1175/JCLI-D-15-0177.1
4. Summary
The 16-yr TRMM radar precipitation features dataset during 1998–2013 is adopted to characterize the large-size warm rain systems in the tropics; they are what we have referred to as the warm organized RPFs. Warm organized systems in the tropics are selected by specifying the precipitation features with minimum infrared brightness temperature warmer than 0°C and rain area greater than 500 km2. Their geographical distribution, properties of precipitation features, and seasonal and diurnal cycles are studied. Several atmospheric fields from ERA-Interim and high-resolution SST are collocated and used to describe the environmental conditions favoring these warm organized rain systems. Finally, the main cause driving the shallow organized RPFs at the southern edge of the eastern Pacific ITCZ is discussed. The primary findings include the following:
Liu and Zipser (2009) have reported some examples of continuous warm rainfall with large area. Based on TRMM RPFs, this study confirms that shallow organized RPFs do exist and they are not uncommon. Mostly they occur over several specific regions, including the eastern Pacific ITCZ, the eastern part of the SPCZ, the south Indian Ocean, the central North Pacific, Hawaii, and several coastal regions between 20°S and 20°N. In contrast with ubiquitous warm but isolated RPFs, warm organized precipitation systems have larger mean rain area varying between 760 and 880 km2 and greater near-surface reflectivity of roughly about 35–40 dBZ. On average, they contribute about 7% of the total warm rain amount over the oceans. The rainfall amounts generated by warm organized systems are greater in winter than in summer.
In comparison to warm isolated systems, warm organized precipitation systems preferably coexist with a drier midtroposphere, a stronger large-scale subsidence at upper levels, an enhanced near-surface moisture convergence, and a stronger low-level convergence.
Although the contribution of warm organized systems to total warm rain is small in a global tropical sense, warm organized systems are shown to have important local impacts over the southern edge of the ITCZ in the eastern Pacific, the SPCZ, and coastal regions. The large-scale atmospheric circulation in the eastern Pacific ITCZ indicates that the prominent shallow meridional circulation tends to be more intense for warm organized RPFs than those for warm isolated RPFs. Namely, warm organized systems are associated with a stronger northerly cross-equatorial return flow from the ITCZ into the Southern Hemisphere at about 800–500 hPa, a southerly inflow into the ITCZ near the surface, and an enhanced ascent in the ITCZ and descending motions to the south. The shallow meridional cell at the north of the warm organized RPFs also appears to be enhanced, compared to isolated ones, but the difference is much less pronounced than the south branch.
Warm organized RPFs in the eastern Pacific are located at the southern edge of warm SST; at the northern side of their convective center, there shows a strong SST meridional gradient. Warm isolated RPFs correspond to relative warmer SST than warm organized RPFs, but the SST gradient northward is significantly weaker, indicating that the SST meridional gradient may be important for the evolution of shallow precipitating systems in the eastern Pacific.
The warm organized RPFs in the tropical eastern Pacific are found to be only at the southern edge of the deep ITCZ core. This is probably related to a meridional asymmetric thermodynamic structure about the equatorial eastern Pacific ITCZ. To the south, there is a stronger shallow meridional circulation that brings more moisture at surface to sustain a large precipitation system. In contrast, the low-level troposphere on the northern side is drier, and there is less moist statistic energy to drive a large-size convective system. In addition, the midlevel dryness and its associated strong subsidence are the keys to maintaining the shallow state of these convective systems.
Acknowledgments
This research was supported by NASA Precipitation Measurement Mission Grant NNX11AG31G under the direction of Dr. Ramesh Kakar. Thanks also go to Dr. Erich Stocker and Patty McCaughey and the rest of the Precipitation Processing System (PPS) team at NASA Goddard Space Flight Center, Greenbelt, MD, for data processing assistance. NOAA high-resolution SST data provided by the NOAA/OAR/ESRL PSD, Boulder, Colorado, from their website (http://www.esrl.noaa.gov/psd/).
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