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2179Bulletin of the American Meteorological Society

1. Traditions

“Stratiform precipitation” is a term often used inmeteorology, yet the term is not defined in the Glos-sary of Meteorology (Huschke 1959), and its mean-ing and usage continually evolve. This article attempts

to clarify the meaning of the term stratiform precipi-tation as it is currently used in the particular contextof tropical meteorology and other regions whereclouds are generated by atmospheric convection. Suchan article would have seemed paradoxical, if not out-right heretical, 30 years ago. Stratiform precipitationwas generally thought to occur almost exclusively withfronts in midlatitude cyclones, where ice particlesgrow predominantly by vapor deposition in a deep,largely buoyantly stable nimbostratus cloud layer, driftdown from upper levels, melt, and fall to the earth’ssurface as raindrops. Under certain cold conditions, theparticles may never melt, and they reach the surface

Stratiform Precipitation inRegions of Convection:

A Meteorological Paradox?

Robert A. Houze Jr.University of Washington, Seattle, Washington

ABSTRACT

It was once generally thought that stratiform precipitation was something occurring primarily, if not exclusively, inmiddle latitudes—in baroclinic cyclones and fronts. Early radar observations in the Tropics, however, showed large ra-dar echoes composed of convective rain alongside stratiform precipitation, with the stratiform echoes covering greatareas and accounting for a large portion of the tropical rainfall. These observations seemed paradoxical, since stratiformprecipitation should not have been occurring in the Tropics, where baroclinic cyclones do not occur. Instead it was fall-ing from convection-generated clouds, generally thought to be too violent to be compatible with the layered, gently set-tling behavior of stratiform precipitation.

In meteorology, convection is a dynamic concept; specifically, it is the rapid, efficient, vigorous overturning of theatmosphere required to neutralize an unstable vertical distribution of moist static energy. Most clouds in the Tropics areconvection-generated cumulonimbus. These cumulonimbus clouds contain an evolving pattern of newer and older pre-cipitation. The young portions of the cumulonimbus are too violent to produce stratiform precipitation. In young, vigor-ous convective regions of the cumulonimbus, precipitation particles increase their mass by collection of cloud water,and the particles fall out in heavy showers, which appear on radar as vertically oriented convective “cells.” In regions ofolder convection, however, the vertical air motions are generally weaker, and the precipitation particles drift downward,with the particles increasing their mass by vapor diffusion. In these regions the radar echoes are stratiform, and typicallythese echoes occur adjacent to regions of younger convective showers. Thus, the stratiform and convective precipita-tion both occur within the same complex of convection-generated cumulonimbus cloud.

The feedbacks of the apparent heat source and moisture sink of tropical cumulonimbus convection to the large-scaledynamics of the atmosphere are distinctly separable by precipitation region. The part of the atmospheric response deriv-ing from the areas of young, vigorous convective cells is two layered, with air converging into the active convection atlow levels and diverging aloft. The older, weaker intermediary and stratiform precipitation areas induce a three-layeredresponse, in which environmental air converges into the weak precipitation area at midlevels and diverges from it atlower and upper levels. If global precipitation data, such as that to be provided by the Tropical Rainfall Measuring Mis-sion, are to be used to validate the heating patterns predicted by climate and general circulation models, algorithms mustbe applied to the precipitation data that will identify the two principal modes of heating, by separating the convectivecomponent of the precipitation from the remainder.

Corresponding author address: Robert A. Houze Jr., Dept. of At-mospheric Sciences, University of Washington, Box 351640,Seattle, WA 98195.E-mail: [email protected] final form: 15 May 1997.©1997 American Meteorological Society

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as snow rather than rain. In both cases, the upward airmotion induces vapor deposition onto the growing iceparticles but is weak enough to allow the particles todrift downward while they grow.

In a baroclinic cyclone, the widespread lifting pro-ducing the growth of the precipitation particles occursin the regions of large-scale warm advection, whichis concentrated in the vicinity of fronts. But stratiformprecipitation also can occur in other large-scale andmesoscale dynamical settings, as long as there is as-cent of saturated air weak enough to allow ice particlesto fall out. In other words, the term stratiform precipi-tation designates a particular set of microphysical pro-cesses leading to the growth and fallout of the precipi-tation within the context of relatively gentle upwardair motion; it does not refer to the specific dynamiccause of the vertical air motions within which the par-ticles form, grow, and fall out.

Since the advent of weather radar in the late 1940s,the term stratiform has been used to describe precipi-tation as it appears in displays of radar data. Stratiformprecipitation is fairly homogeneous in the horizontal,giving it a layered structure in vertical cross sectionsof radar reflectivity. In particular, it often exhibits apronounced layer of high reflectivity called the “brightband,” marking the layer in which the downward settlingice particles are melting (Battan 1973; Houze 1993).

Stratiform radar echoes contrast sharply with radarechoes from “convective precipitation,” which appearson radar as “cells,” another loosely defined term, ap-parently coined by Byers and Braham (1949) in theirclassic report on the Thunderstorm Project in Ohio andFlorida. Cells are horizontally localized patches orcores of intense radar reflectivity. In a vertical crosssection, a cell is a tall, thin column of high reflectivity.The orthogonality of its structure to that of abrightband echo often makes the convective “cell”quite recognizably distinct from “stratiform” precipita-tion on radar, and radar meteorologists have developeda tradition of speaking of radar echoes as convective orstratiform, according to whether they form patterns ofvertically oriented intense cores or horizontal layers.

2. Radar observations in tropical fieldexperiments: Emergence of aparadox

Because of its ubiquitous presence in barocliniccyclones and fronts, and because of the dearth of me-teorological observations in the Tropics up to the sec-

ond half of this century, stratiform precipitation wasthought to occur primarily, if not exclusively, inmiddle latitudes. As late as 1979, Herbert Riehl re-ferred to tropical rain as “a cumulus regime, in con-trast to the climates with stratus precipitation in higherlatitudes.” Riehl was correct in characterizing theTropics as a “cumulus regime”; precipitating cloudsin the Tropics are entirely convection-generated cumu-lonimbi, and baroclinic cyclones and fronts, indeed,do not occur in the Tropics. However, his statementdoes not recognize that stratiform precipitation is notrestricted to higher-latitude clouds. It may also occurwithin the “cumulus regime.”

In 1972, scientists aboard a Soviet research shipoperating in the intertropical convergence zone (ITCZ)of the eastern tropical Atlantic took photographs of aradar display showing strong brightband echo(Shupiatsky et al. 1975, 1976a,b). In 1974, the GlobalAtmospheric Research Programme Atlantic TropicalExperiment (GATE) obtained more extensive radarobservations in this region. One of the most signifi-cant findings of GATE was that the stratiform radarechoes, seen earlier by the Soviet scientists, had verystrong bright bands and covered large areas (Houze1975, 1977; Leary and Houze 1979a,b). About 40%of the rain falling on the ocean surface in GATE wasstratiform (Cheng and Houze 1979; Leary 1984).These observations showed definitively that both con-vective and stratiform radar echoes occurred in a re-gime in which the clouds are generated entirely byatmospheric convection.

Since GATE, radar studies have confirmed, manytimes over, that a large component (~20%–50%) oftropical precipitation exhibits stratiform radar-echostructure [see Houze (1989) for a summary]. More-over, the stratiform echoes usually coexist with cellswithin a “mesoscale” precipitation area, in which therain covers contiguously a region 10–1000 km inhorizontal dimension. Higher rainfall rates in thesemesoscale precipitation areas are concentrated in thecells, but lighter stratiform rain covers most of themesoscale area. A portion of the area is covered by echoof intermediate intensity, which may be either “convec-tive” or “stratiform,” depending on its vertical structure.

In addition to the purely tropical regimes, in whichthere is no chance of a baroclinic contribution to theprecipitation processes, there have now been manystudies recognizing the large presence of stratiformprecipitation in convection-generated clouds and pre-cipitation in midlatitude regimes (e.g., Rutledge andMacGorman 1988; Houze et al. 1990). The discussions


in this paper apply to both midlatitude and tropicalconvection.

3. Consistent terminology: The purposeof this paper

The fact that tropical precipitation, when observedby radar, has a substantial stratiform component hasled to a terminology that seems contradictory, sincethe tropical atmosphere is clearly a region of convec-tion. The term convection refers to the overturning ofa fluid under gravity, and in the Tropics, nearly allprecipitation (except part of that associated with tropi-cal cyclones and possibly some orographic rain) is theproduct of atmospheric convection. However, the termstratiform, as defined above, is an adjective describ-ing a set of microphysical processes that occur inweaker vertical air motions. Similarly, the term “con-vective” describes a radar echo from precipitationgrowing by the cloud-microphysical processes thatattend the strong vertical air motions of young, activeconvection. Since stratiform radar echoes may be seen

in an atmosphere of older convection, the terms strati-form (an adjective) and convection (a noun) are notmutually exclusive; stratiform describes the nature ofsome of the precipitation growth processes occurringin the region of convection, while the term convectionalludes to the fluid dynamical origin of the air motions.

Henceforth, this paper, which is about tropical andother convection-generated precipitation, will use thenoun convection to refer to an overturning fluid andthe adjective convective to describe the precipitation(or radar echo) associated with young, active convec-tion. The adjective stratiform will refer to precipita-tion occurring in older, less active convection andpossessing radar echoes that have weak horizontalgradients and/or a bright band. Describing the radarechoes as being either convective or stratiform thusimplies that the echoes depict precipitation in regionsof stronger or weaker air motions, respectively, whileat the same time recognizing that both the weaker andstronger air motions may arise from the process ofconvection. Convection can lead to cumulonimbusclouds, whose precipitation is partly convective andpartly stratiform.1 Figures 1 and 2 illustrate, in ideal-

FIG. 1. Conceptual models of vertical cross sections through (a)–(c) young, vigorous precipitating convection and (d)–(f) old convection.

1The stratiform precipitation could be considered to be falling from nimbostratus cumulonimbogenitus (precipitating stratiform cloudarising from cumulonimbus). Since such nimbostratus is usually a cloud deck currently or previously extruding out of the cumulo-nimbus, it seems artificial to call it a separate cloud. The author prefers to call the whole, single precipitating cloud entity a cumulonimbus.

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ized form, the terminology of this paper. The remain-der of this paper delves further into this idealized pic-ture in order to reach a deeper understanding of con-vection-generated precipitation and how to observe,describe, and interpret it.

4. What is convection?

To understand how both convective and stratiformprecipitation can occur in response to convection, weshould first review what we mean by convection in thecontext of cloud-formation processes. We say convec-tion occurs when a fluid under gravity is heated from

below or cooled from the top at such a rate that mo-lecular diffusion cannot redistribute the modifieddensity field fast enough so as to maintain equilibrium.The fluid layer then becomes buoyantly unstable andoverturns macroscopically to stabilize the stratifica-tion of the density. One of the simplest examples isBenard convection in an incompressible fluid layer,where the fluid overturns by means of a field of geo-metrically simple vertically circulating elements. Theatmospheric convection that produces cumulus andcumulonimbus clouds has the additional complica-tions that air is a compressible fluid and contains wa-ter vapor, so overturning acts to neutralize the stratifi-cation of moist-static energy. Figure 1a shows atmo-spheric convection as a pattern of idealized verticallycirculating elements, which are geometrically sym-metric (like Benard cells) in the drawing, although inreality, the shapes of the overturning elements are highlymodified by wind shear and by the interactions of thecloud and precipitation fields with the air motions.Asymmetries in the vertical structures of the cells donot affect the present discussion, which aims only todistinguish the terms “convective” and “stratiform” indescribing precipitation. At the most basic level, itserves to keep to the symmetric idealization in Fig. 1a.

In atmospheric convection, water may condenseand fall out. A saturated layer of air that is convectivelyoverturning and precipitating, therefore, differs froma layer of Benard convection in that there must be anet upward transfer of air between the upper and lowerboundaries of the fluid in order to account for the netcondensation of water in the layer (as long as the rateof rise of temperature is less than the heating rate). Theupdrafts of the overturning cells must, therefore, trans-port more mass on average than do the downdrafts.Figures 3a and 3b summarize the effects of verticalmass transport in a region of active convective cells:net upward mass transport produces net latent heatingat all levels (Fig. 3a), and mass continuity requires nethorizontal convergence at low levels and divergencein the upper troposphere (Fig. 3b, open arrows in Fig.1a). In the real atmosphere, a region of young, vigor-ous convection may not be completely saturated in itsdowndraft regions, but measurements confirm that thenet mass divergence profiles are like that in Fig. 3b(e.g., Mapes and Houze 1995).

5. What is convective precipitation?

In deep precipitating atmospheric convection, the

FIG. 2. Idealization of a horizontal map of radar reflectivity (a)divided into convective and (b) stratiform regions.


vertical velocities in the cores of the updrafts are sev-eral meters per second or greater. These strong updraftscondense vapor rapidly, producing large concentra-tions of cloud liquid water. Houghton (1968) pointedout that, in the presence of such strong updrafts, thedominant growth mechanism for precipitation par-ticles is the collection of this cloud water by largerdrops and/or ice particles sweeping out the cloud wa-ter in their fall paths. For growing water drops, thisprocess is termed “coalescence”; in the case of collec-tion by ice particles, it is known as “riming.” More-over, the strong updrafts allow for the larger particlesto grow for a long period of time because they are car-ried upward relative to the earth, even though they arefalling relative to the smaller cloud water droplets.This upward advection of the growing particles in-creases their residence time in cloud and thus theiropportunities to collect cloud droplets. Figure 4 illus-trates how, as a parcel of updraft air rises, the grow-ing particles within it move up until they become largeenough to fall relative to the air. At each successiveheight, some particles fall out of the parcel, while theremainder, which are not as heavy, are spread out lat-

erally over an increasingly greater area by the diver-gent air flow as the rising, buoyant air parcel expands.

Since the bulk of the precipitation mass falls outwithin a few kilometers of the updraft centers, the ra-dar reflectivity pattern associated with the convectiveair motions in Fig. 1a is a set of concentrated peaks ofreflectivity, as shown in Fig. 1c. In plan view, the con-vective region appears in the radar echo as a field oflocalized reflectivity maxima (Fig. 2).

6. Stratiform precipitation in oldconvection

When atmospheric convection is young and vigor-ous, the precipitation occurs in cells, as indicated con-ceptually in Fig. 1a. However, as a region of convec-tive cells weakens, the precipitation associated withthe cells takes on a layered structure resembling thatfound in extratropical cyclonic precipitation. The up-ward vertical air motions in the region of weakenedcells are strong enough to allow precipitation particlesto grow by vapor diffusion but too weak to supporthigh concentrations of cloud liquid water; therefore,growth by riming is not so effective as when the up-drafts are stronger. In this respect, then, the region ofweakened cells in a tropical cloud system resemblesan extratropical cyclone.

Thus, when we divide the precipitation area of aconvection-generated cloud system into convectiveand stratiform components (Fig. 2), we imply that the

FIG. 3. Characteristic profiles of latent heating and horizontalmass divergence in convective and stratiform regions of tropicalprecipitation.

FIG. 4. Conceptual model of an updraft behaving like a buoyantbubble with different entrainment rates. Dots indicate hydro-meteors suspended by updraft; downward-pointing arrowsindicate particles heavy enough to fall through updraft; horizontalarrows indicate lateral spreading of bubble. Open arrows representthe vector field of the buoyancy pressure gradient force (as in Fig.7.1 of Houze 1993). [From Yuter and Houze (1995c).]

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stratiform region is dominated by weaker air motions.These vertical air motions could be horizontally uni-form across the stratiform region; however, they areusually variable in the horizontal. The requirementsare only that there be net upward mass transport (toallow net condensation of vapor) in the cloud layer inwhich particles are growing on average and that theensemble of vertical air motions in the cloud layercontain few if any upward velocities > 1 m s−1 (so thatcloud liquid water production, and hence riming, isminimal and precipitation-sized particles are not ad-vected upward, so as to overly disrupt the layeredreflectivity pattern).

In a deep precipitating tropical cloud system, the“stratiform” precipitation region is typically a regionof older convection (Fig. 1d). It is well known that inthe older portions of a cumulonimbus, the lower tro-posphere is dominated by downdrafts, while the up-per levels are dominated by weak updrafts. The upperlevels are not wholly occupied by updrafts, nor are thelower levels wholly occupied by downdrafts. In theolder part of the precipitation area of a Florida cumu-lonimbus, the vertical mass transport at upper levelswas net upward (Fig. 5b), such that there was netgrowth of precipitation particles. However, the netupward mass transport at upper levels was the com-bined effect of an ensemble of vertical velocities, in-cluding some negative (downward) values, as shownby the distribution of the mass flux as a function ofthe vertical air velocity and height in Fig. 5a. How-ever, updrafts < 2 m s−1 accounted for most of the up-ward mass transport; that is, nearly all of the volumeof the storm at upper levels was occupied by air mov-

ing upward too slowly to overcome the downward fallvelocities of precipitation particles. At these verticalair velocities, the condensed water is transferred to thegrowing ice particles almost entirely by vapor diffu-sion (Rutledge and Houze 1987), and the particles driftdownward, as suggested by Fig. 1e. Below the 0°Clevel, the older precipitation regions are dominated bynet downward motion (Houze 1989). Figure 5 showsthat updrafts, mostly < 1.5 m s−1 in intensity, are in-terspersed among more predominantly downward airvelocities. Zipser and LeMone (1980) and LeMoneand Zipser (1980) described how research aircraft inGATE encountered both up- and downdrafts in osten-sibly stratiform parts of tropical cumulonimbus.

When growing ice particles drift downward withinthe mid- to upper altitudes of a volume of air within aregion of old convection such as that portrayed in Figs.1d–f, distinct layers become evident via contrastingmicrophysical processes (Fig. 1e). At upper levelsvapor diffusional growth is prevalent, and it is too coldfor aggregation to occur. Empirical evidence indicatesthat aggregation occurs at ambient temperatures of 0°to ∼ −15°C with the highest probability of aggrega-tion in the temperature range of about 0 to ∼ −5°C(Hobbs 1974, 641). Houze and Churchill (1987) con-firmed this layering in tropical convection: they foundlarge aggregates in this temperature layer among theparticle images collected by aircraft in the stratiformregions of tropical convection over the Bay of Bengalin the Global Atmospheric Research Programme’sMonsoon Experiment.

Displays of radar data enhance the layered appear-ance of stratiform precipitation because the reflectivity

FIG. 5. (a) Vertical mass transport distribution by vertical velocity and (b) mass flux in a Florida cumulonimbus. Units of contoursare [mass transport/(dw dz)], where dw = 1 m s−1 and dz = 0.4 km. Contours are at intervals of 25 × 106 kg s−1. [Adapted from Yuterand Houze (1995c).]


of the melting snow is especially high for two reasons:1) some of the melting particles are very large aggre-gates of ice crystals, and reflectivity is proportional tothe sixth power of the particle dimension, and 2) theindex of refraction of melting snowflakes is ~5 timesgreater than nonmelting ice. In addition, when the iceparticles melt, their fall speeds increase by a factor ofabout 5, and the particles evacuate the melting layerrapidly in a vertically divergent fashion, producing adecreased concentration of particles below the melt-ing layer. These factors combine to produce a pro-nounced bright band of enhanced reflectivity in a shal-low layer centered just below the 0°C level (Fig. 1f).

The bright band is an unambiguous indicator of thepresence of stratiform precipitation. However, theabsence of a bright band does not imply the absenceof stratiform precipitation structure. A bright bandwill not be observed unless the vertical resolution ofthe radar is sufficiently fine to delineate the band.Hence, the bright band is a property of the radar data;it is not a unique measure of the magnitude of themelting. A strong bright band only will appear if someof the melting particles are in the form of large aggre-gates. Temperature, ambient humidity, and crystalstructures must all meet certain criteria for aggrega-tion to occur. Aggregation of precipitating ice particlesto form snowflakes does not add any mass to the pre-cipitation, nor do the particles fall any faster once theyhave aggregated. Just as much latent heat is consumedin the melting, and the vertical flux of precipitationmass remains the same, regardless of whether or notthe precipitating ice particles happen to aggregate toform large highly reflective particles and hence abright band. Braun and Houze (1995) showed that themelting layer of a midlatitude mesoscale convectivesystem was much more horizontally extensive than theradar bright band, which was present only in the partof the melting layer in which melting ice particles hadaggregated to form very large snowflakes.

The presence of a bright band thus indicates thepresence of a particular type of stratiform precipitation,namely that in which large aggregate snowflakes areamong the melting particles. Since the ability to de-tect a bright band is a function of the characteristicsof the radar rather than the storm, further caution isrequired, since the absence of a bright band does notindicate the absence of a layer of melting aggregates.Typically, a radar detects a bright band only at closerange because the radar beam broadens with distancefrom the antenna and the melting layer is only ~0.5 kmor less in depth.

The idealized bright band sketched in Fig. 1f ishorizontally inhomogeneous, broken into severalpatches, and in one patch “fallstreaks” extend down-ward from the melting layer. Stratiform precipitationin tropical cumulonimbus usually occurs in regionswhere previously intense convection (like that de-picted in Figs. 1a–c) has weakened. Since the previ-ous convection was concentrated in cells, the result-ing stratiform structure remains somewhat patchy,with highest reflectivities in the regions where convec-tive cores once were vigorous; inspection of the high-resolution airborne radar data obtained in TOGACOARE2 shows fallstreaks consistently in the loca-tions of previously active convective cells (e.g., Yuterand Houze 1997a). The fallstreaks appear to occurwhere the remnant precipitation cores have not com-pletely lost their identities, and since they are no longerpart of an active convective cell, the remnant showersare distorted into tilted, bent fallstreaks by the ambi-ent wind shear.

An alternative explanation for the fallstreaks is thatan unstable layer produced by the latent cooling of theair by the melting leads to the formation of small, shal-low convective cells in the melting layer. In this case,the fallstreaks would be similar to the “generatingcells” seen in extratropical cyclonic precipitation(Marshall 1953). It is possible that the fallstreaks intropical convection are a combination of remnant pre-cipitation cores and overturning induced by meltingcooling.

7. Dynamic implications of thestratiform component of tropicalprecipitation

Although vertical air motions are relatively weakin a stratiform region, the area covered by the strati-form precipitation can be much larger than that occu-pied by the active convective cells. A large portion ofthe time- and space-integrated vertical mass transportof the convection thus occurs in the older convection,that is, in the stratiform regions of cumulonimbus. Forthis reason, the stratiform regions of tropical convec-tion have important dynamic implications.

2The Tropical Ocean Global Atmosphere Coupled Ocean–Atmo-sphere Response Experiment (TOGA COARE) was a large fieldexperiment employing aircraft and ships to study the behavior ofthe atmosphere and ocean over the western tropical Pacific Ocean.The experiment ran from 1 November 1992 to 28 February 1993.

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The large-scale environment responds to a convec-tive system in a manner analogous to the way thewater responds to a rock dropping into a pond. Thewater has to make way for the negatively buoyant rockaccelerating downward through the fluid environment.The fluid adjusts to the sudden displacement of thewater by a spectrum of gravity waves, which rippleaway from the spot where the rock falls in the water.

In a similar way, a buoyant parcel of air pushesambient air out of its path, and the surrounding atmo-sphere adjusts by sending out a spectrum of gravitywaves, whose propagation speeds are proportionalto their vertical wavelengths (Bretherton andSmolarkiewicz 1989; Mapes 1993). In a region of pre-cipitating convection, there is a net positive buoyancyproduced by the latent heat gained by air when pre-cipitation falls out. The surrounding atmosphere ad-justs to this generation of buoyancy by a spectrum ofgravity waves (or more accurately bores) that have thenet effect of displacing mass outside the convectiveregion downward (Mapes 1993; Mapes and Houze1995). The open arrows in Fig. 1a show the net resultof this compensating downward displacement as a nethorizontal transport of air from the environment intothe active convective region (mass convergence) atlower levels and out of the region (mass divergence)at upper levels.

By the time the convective region is older, the netvertical mass transport in the lower levels of the pre-cipitation region is downward (the negatively buoy-ant downdrafts begin to dominate over the updrafts).Net cooling occurs as a result of melting and evapo-ration at mid- to low levels during this stage (Fig. 3c),and horizontal divergence of air out of the storm mustcompensate the negative buoyancy in the lower tro-posphere of the rain areas (Fig. 3d, and the open ar-rows at low levels in Fig. 1d). Gravity waves that re-arrange the environmental mass field around the strati-form rain areas must produce a net upward displace-ment of the environment to compensate the net hori-zontal mass divergence out of the old convective (i.e.,stratiform) region at low levels.

The open arrows in Fig. 1d also show a net hori-zontal transport of air into the old convective region(mass convergence) at midlevels. This midlevel con-vergence compensates both the (net) positive buoy-ancy of the upper levels of the stratiform region andnet negative buoyancy of the lower levels.

The most direct way to measure the adjustment ofthe large-scale environment to a region of convectionis to measure the horizontal mass divergence into and

out of the region. Accurate measurement of the out-ward normal component of the wind along the bound-ary of the volume of atmosphere containing the con-vection gives us this divergence, and in tropical fieldexperiments, airborne Doppler radar data and rawin-sondes have been used for this purpose. In TOGACOARE, aircraft Doppler radar provided 143 profilesof divergence around regions of precipitation. An ex-ample of the average of the profiles obtained on a flightsampling primarily convective precipitation showedconvergence at low levels and divergence at upper lev-els (Fig. 6a), consistent with the open arrows in Fig.1a. A flight sampling primarily old convection, indi-cated by the stratiform character of the radar echo,indicated strong convergence at midlevels (Fig. 6b),consistent with the open arrows in Fig. 1d. Figure 7shows the grand mean profile of divergence, obtainedby averaging all 143 samples obtained on flights inTOGA COARE. It represents the combination of ac-tive and old (stratiform) convection sampled on all theflights, over a 4-month period. The solid curve in Fig.8 shows the profile of divergence obtained by averag-ing TOGA COARE rawinsonde data obtained aroundregions of satellite-observed active deep convection.It is consistent with the profile of divergence from theairborne Doppler radar (compare the profile in Fig. 7with the solid curve in Fig. 8).

To determine how the large-scale atmosphere ad-justs to a disturbance of its mass field by a region ofdeep convection, Mapes and Houze (1995) used themeasured divergence profiles in Figs. 7 and 8 as inputto a linear spectral primitive-equation model. In thismodel the divergence profile serves as a mass distur-bance for the large-scale flow. The time-dependentflow predicted by the model, using the observed massdivergence profiles as input, indicates the modes ofmotion that develop when the large-scale atmosphereis perturbed by a divergence profile similar to thatmeasured in the vicinity of the deep convection occur-ring in TOGA COARE. The calculation applies theobserved divergence profile impulsively over a 140-km diameter region (a typical size of a deep convec-tive precipitation area in TOGA COARE).

Two distinct modes dominate the model’s large-scale response to the convection. The solutions of theequations are decomposed into vertical wavelengthcomponents, which correspond to the vertical wave-length components of the spectrally decomposed ob-served profile of divergence. The wave speed is di-rectly proportional to the vertical wavelength, and Fig.9a shows the magnitudes of the spectral coefficients


of the divergence for each wave speed. The bigger thecoefficient, the more that component of the divergenceprofiles accounts for the large-scale response to the

convection. Figure 9a shows that the motions aredominated by bands of wave speeds centered on 52and 23 m s−1.

FIG. 6. Mean profiles of divergence measured by airborne Doppler radar in tropical precipitation over the western tropical Pacificduring TOGA COARE: (a) from a flight in which most of the sampled precipitation was in a vigorously convective state (average of13 samples obtained 15 December 1992); (b) from a flight in which most of the sampled precipitation was in a stratiform state (averageof 15 samples obtained on 6 November 1992). All samples are for a region 30 km in diameter. The three lines in each plot are threeindependent radar measures for negligibly different geometry. [From Mapes and Houze (1995).]

FIG. 8. Divergence measured by rawinsondes when extensive,deep convection-generated cloud systems affected the intensiveflux array (centered near 2°S, 156°E) over the western tropicalPacific during TOGA COARE. Sixteen cases were used tocompute the net divergence (solid curve). The components of thenet divergence constituted by the 52 and 23 m s−1 gravity waveresponses to the mass disturbance are shown by the dashed anddashed–dotted curves, respectively. [Derived from calculations ofMapes and Houze (1995) and provided by B. Mapes.]

FIG. 7. Mean of all divergence profiles measured by airborneDoppler radar in tropical precipitation over the western tropicalPacific during TOGA COARE. The average is based on 143samples, obtained on 10 different aircraft missions. [From Mapesand Houze (1995).]

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Figure 8 shows the computed 52 and 23 m s−1 com-ponents of the divergence profile in comparison to thetotal observed profile of divergence. The 52 m s−1 com-ponent has a two-layer structure, consisting of conver-gence at low levels and divergence at upper levels,while the 23 m s−1 component has a three-layer struc-ture, consisting of convergence at midlevels and di-vergence at upper and lower levels. Evidently, thesetwo dominant modes of the observed divergence pro-

file correspond respectively to the convective andstratiform portions of the precipitation pattern. Addedtogether, the convective (52 m s−1) and stratiform (23m s−1) curves in Fig. 8 account for nearly the entireobserved divergence (solid curve).

The atmospheric adjustment to the active convec-tive region is a fast-moving gravity wave “bore” thatdisplaces environmental mass downward, to mass-balance the convergent inflow to the convective regionat low levels and the divergent outflow aloft. Since thisbore moves away at 52 m s−1 (over 100 mph), the ad-justment is very quick. The adjustment to the strati-form region’s older convection is less than half as fast,and this bore displaces mass upward in the lower tro-posphere and downward in the upward troposphere toadjust for the stratiform region’s convergence atmidlevels and divergence at upper and lower levels.

Figure 9b shows the effect of the fast (convective)and slow (stratiform) bores 6 h after the impulsive startof the heating. The convective bore has produced amass-compensating downward displacement throughthe depth of the troposphere at 600–1300 km from thecenter of the disturbance. This compensating down-ward motion suppresses convection at these distances.At 150–600 km from the source, the stratiform bore,which has moved out only about half as far as the con-vective bore, has produced a net upward displacementof mass in the lower troposphere and a net downwarddisplacement aloft. An important implication is that adeep convective tropical disturbance may actuallyencourage new convection in its immediate vicinityvia the compensating upward motion in the lower tro-posphere. This result led Mapes (1993) to dub tropi-cal convection “gregarious.”

8. Convective and stratiformprecipitation: A dichotomy or theends of a spectrum?

The results of Mapes and Houze (1995) in Figs. 8and 9 suggest that there is no reason, from the view-point of large-scale dynamics, to distinguish catego-ries of air motion in a tropical precipitation regionother than the categories convective and stratiform—that is, a dichotomous structure of convection. The im-pression formed from looking at radar-echo maps oftropical precipitation, however, is that a significantportion of the precipitation field is difficult to charac-terize as convective or stratiform. In an earlier paper,Mapes and Houze (1993) classified one-third of the

FIG. 9. Results of a linear spectral model of a large-scaletropical atmosphere perturbed by a mass divergence profile. Thesymbol δ

d stands for the “diabatic divergence,” which is the mass

disturbance produced by the net heating of a mesoscaleprecipitating cloud system embedded in the large-scale flow. Theobserved divergence profile used for δ

d is based on airborne

Doppler radar and sounding observations obtained over thewestern tropical Pacific in TOGA COARE in regions containinglarge cumulonimbus cloud systems. This observed divergenceprofile is applied impulsively at the initial time t over a region 140km in horizontal dimension with a horizontal cos2 profile. Themodel computes the response of the large-scale environment tothe cloud disturbance at all times after the impulsive massdisturbance. The response is a spectrum of gravity waves withvarying vertical wavelengths. Each wavelength corresponds to awave speed c

n. The symbol δ

dn in (a) represents the spectral

coefficient (magnitude) of the response at each wave speed.Temperature perturbation T shown by the contours in (b) isproduced by the fast (convective) and slow (stratiform) gravity-wave bores 6 h after the impulsive start (time t) of the heating.[Adapted from Mapes and Houze (1995).]


radar echoes observed with the WP-3D radars over thetropical ocean north of Australia as “intermediary,”because the echo structure appeared to be neither ob-viously convective nor obviously stratiform, and be-cause these echoes appeared to be in the process ofconverting from convective to stratiform structure.Mapes and Houze’s (1993) findings thus suggest thata spectrum of structures exists between the convectiveand stratiform extremes, an apparent contradiction tothe 1995 results.

The key to resolving this apparent contradiction isto keep in mind that radar-echo structure is not asingle-valued function of the vertical air motions. Ifthe intermediary radar-echo structures represented asignificantly different kinematic category, the mass-divergence spectrum (Fig. 9a) would not be so decid-edly bimodal. Each instantaneous two-dimensionalcross section of radar data (horizontal or vertical) con-tains only a selected subset of the data in a three-di-mensional volume. A natural tendency is to draw crosssections through extrema, which give the cross sec-tions a rougher (more “convective”) appearance. Amore representative and complete way of viewing theechoes is to combine all the data in the three-dimen-sional volume of space scanned by a radar in a singleplot and inquire about the statistics of the radar ech-oes in that volume. Yuter and Houze (1995a–c) presentthe radar data from a three-dimensional volume ofspace as a joint probability distribution, which theyrefer to as a “contoured frequency by altitude dia-gram,” or CFAD. The CFAD of radar reflectivity con-sists of contours showing the range and relative nor-malized frequency of occurrence of radar reflectivityvalues as a function of reflectivity and height aboveground.

Figure 10a shows a CFAD of the radar reflectivityvalues observed in a convective precipitation regionof a Florida cumulonimbus. The distribution ofreflectivities at all altitudes is broad, and even at highaltitudes a few rather high reflectivities occur, indicat-ing the presence of graupel or small hail, which is con-sistent with growth by riming in strong updrafts. Incontrast, the stratiform region CFAD of a squall linein Kansas (Fig. 10c)3 shows a narrower distribution ofrelatively uniform values of reflectivity at all altitudes,and at upper levels, the mode is a much lower

reflectivity, consistent with a dominance of vapordeposition as the primary microphysical growth modeof the particles. The mode of the distribution increasesas the altitude decreases, indicating the microphysi-cal growth and aggregation of the falling precipitationparticles as they approach the melting level. Just be-low the 0°C level (~4 km), the mode jumps to a highvalue, corresponding to the typical reflectivity of themelting aggregates falling through the layer.

As the storm within the three-dimensional volumescanned by a radar ages, the CFAD of reflectivitychanges from a more convective to a more stratiformstructure. Yuter and Houze (1995b) found thatchangeover occurred quickly, even while the stormstill contained a few intense up- and downdrafts andbefore the storm appeared obviously stratiform in in-dividual vertical cross sections of reflectivity. TheCFAD in Fig. 10b is for a time in the mid- to late stagesof the same storm whose early convective stage is rep-resented in Fig. 10a. Although the storm still containeda wide range of vertical velocities, the reflectivity fieldwas already statistically more stratiform, with a nar-rowing of the distribution. This result of Yuter andHouze (1995b) suggests that the intermediary radar-echo structures identified by Mapes and Houze (1993)would have been categorized as stratiform had theybeen viewed as CFADs.

Yuter and Houze’s (1995b,c) results further indi-cate that even while some strong up- and downdraftswere still active, and the radar echo pattern seen inhorizontal and vertical cross sections was nonuniform,most of the mass flux of the storm was in updrafts in-sufficiently strong to prevent ice particles from fall-ing or to produce much liquid water to be collected bythe falling ice particles. In a statistical sense, then, mostof the volume of the storm was occupied by ice par-ticles growing primarily by vapor deposition and fall-ing relative to the ground; that is, the stratiform pre-cipitation process dominated in most of the volume ofthe storm even though some spots within the storm stillhad strong updrafts.

Mapes and Houze (1993) used a large set of air-borne Doppler-radar measurements to determine thevertical profiles of divergence in the convective, in-termediary, and stratiform radar echo structures theyobserved over the tropical ocean north of Australia

3Figure 10c is not from the same case as (a) and (b), because the observations in the Florida case illustrated in (a) and (b) ended beforethe fully developed stratiform region could be observed. The convective region of the Kansas case illustrated in (c) could not be usedfor (a) and (b) because the spatial resolution of the radar data in the Kansas case was too coarse to show the detailed structure of theconvective region. Hence, Fig. 10 had to be pieced together from the two cases.

2190 Vol. 78, No. 10, October 1997

(Fig. 11). Around intermediary echo structures, theyfound divergence in low to midlevels overlain by alayer of convergence (Fig. 11b).4 This behavior is fur-ther evidence that the intermediary echo regions were

kinematically more like stratiform than convectiveprecipitation regions. However, the layer of conver-gence was at a somewhat higher altitude than in thestratiform regions (cf. Figs. 11b,c). This differencemay have been a result of sampling (the data in Fig.11b were not necessarily obtained in the same stormsas the data in Fig. 11c), or there may be a slight dif-ference between intermediary and stratiform dynamics.

Figure 12 suggests possible relationships betweenthe pattern of up- and downdrafts and the correspond-ing divergence profiles during the various stages ofconvection. In regions dominated by active convection(Fig. 1a), strongly buoyant, undiluted elements rise,and mass continuity requires that the buoyant elementsbe surrounded by a pressure-gradient force field, whichpushes air out of the path of the top of the cell and intoward the base of the buoyant updraft in the lower

FIG. 10. Examples of CFADS of reflectivity for (a) region of young, vigorous convection; (b) a region of intermediary convection;and (c) a stratiform precipitation region. Units of contours are (frequency of occurrence, in percent of observations at a given altitude)/(dz ddBZ), where dz = 0.4 km and ddBZ = 5. Contours are drawn at intervals of 2.5% dBZ−1 km−1. The 5% dBZ−1 km−1 contour ishighlighted. [Adapted from Yuter and Houze (1995b).]

4All the profiles in Fig. 11 should have a layer of divergence atupper levels. The sensitivity of the radar used to obtain the diver-gence profiles in Fig. 11 was, however, not sufficient to observethe low intensity of the reflectivity at these high levels. Other datashowed that the actual cloud tops were several kilometers abovethe 200-mb level (~12 km), which is the highest level at whichthe divergence could be reliably measured by the radar. All themean profiles showed a net convergence below 12 km. On thelinear pressure scale used in Fig. 11, the areas under the diver-gence profiles are proportional to mass divergence. Since mass con-vergence and divergence must balance over a full vertical column,the net convergence seen in Fig. 11 implies that the higher levelsmust have been characterized by net divergence in all three cases.


troposphere. In a region containing one or more active,deep convective elements, both positively buoyantupdrafts and negatively buoyant downdrafts arepresent in the lower troposphere (Fig. 12a). Over theregion of convection, the upward vertical mass flux ofthe updrafts dominates over the downward mass flux

of the downdrafts. In the boundary layer, some of theconvergence into the base of the buoyant updrafts iscanceled, in an area-averaged sense, by divergenceassociated with downdrafts. Thus, the maximum con-vergence into the convective region of a mesoscaleconvective system is usually found at some heightabove the boundary layer. Even so, the convective-region divergence profile is two layered, with thelower troposphere characterized by net convergenceand the upper troposphere by net divergence.

As a strongly buoyant element rises in a cumulon-imbus cloud, it typically gets cut off from its supplyof high moist-static energy air from below. When itreaches the stable tropopause, it spreads out, collaps-ing into a flattened buoyant element (Lilly 1988). Theconvergence produced by the pressure-gradient forceat the base of the buoyant element is then elevated into

FIG. 11. Mean divergence observed by airborne Doppler radar in tropical precipitating cloud systems over the ocean north ofAustralia. [Adapted from Mapes and Houze (1993).]

FIG. 12. Conceptual models explaining the vertical profiles ofdivergence in regions of precipitation dominated by (a) young,vigorous convection; (b) intermediary convection; and (c)stratiform precipitation. In (a) and (b), the rectangles encloseregions of positive (+) and negative buoyancy (−) associated withup- and downdrafts, respectively. The arrowheads indicate thedirection of the pressure-gradient force field induced by thebuoyancy field [see Fig. 7.1 of Houze (1993).] The divergenceprofiles on the right-hand sides of (a) and (b) reflect the net effectof these force fields. In (c), no attempt is made to identifyindividual buoyant or negatively buoyant elements, as theprecipitation (outlined by the radar-echo boundary) contains theremains of previously vigorous buoyant and negatively buoyantelements, which intermingle such that the upper levels aregenerally positively buoyant, and the lower levels are generallynegatively buoyant. The heavy shading in (c) indicates the brightband just below the 0°C level and the remnants of old cellsdistorted into fall streaks by the wind shear.

2192 Vol. 78, No. 10, October 1997

the upper troposphere. Figure 12b suggests that anintermediary echo region, of the type analyzed byMapes and Houze (1993), is dominated at upper lev-els by buoyant elements at this stage of developmentand at lower levels by negatively buoyant downdrafts.The net profile of divergence in a region of interme-diary echo structure would then look like that on theright-hand side of Fig. 12b, with a layer of maximumconvergence in the upper troposphere correspondingto the convergence at the base of the elevated, hori-zontally spreading buoyant elements. This hypotheti-cal profile of divergence is consistent with the observeddivergence profile in intermediary regions (Fig. 11b).

In the stratiform region, negatively buoyantdowndrafts dominate the lower troposphere, whileweakened, but still positively buoyant updrafts, domi-nate the upper troposphere, and a strong melting layerproduces a radar bright band in the middle troposphere(Fig. 1b). By this time, a lot of individual buoyant el-ements have risen into the upper troposphere, reachedtheir level of zero buoyancy, and spread laterally. Somereach the tropopause, but others entrain and come torest at lower levels. The net result is that the whole mid-to upper troposphere above the melting layer is filledwith intermingled old buoyant elements (Fig. 12c). Atthis stage, no strongly buoyant elements dominate themean buoyancy; rather, the conglomeration of old el-ements dominates the picture. So Fig. 12c shows themid- to upper levels as generally positively buoyantbut does not attempt to identify individual elements.

A general region of negative buoyancy dominatesthe stratiform region from the 0°C level down to theearth’s surface (Fig. 12c). This region of negativebuoyancy, in a mature stratiform region, is a combi-nation of old negatively buoyant convective elements,cooling of the air in the melting layer, and entrainmentof low moist-static-energy air from midlevels of theenvironment. The pressure force field required by theresidual positive buoyancy aloft and the negative buoy-ancy of both the downdrafts and the melting layer leadto a maximum of convergence just above the 0°C level.

By comparing the divergence profiles in Figs. 12band 12c, we infer that they differ only in the altitudeof the maximum convergence into the region, whichis slightly higher in the intermediary case. Thus, whileon the order of one-third of the reflectivity patternshave an intermediary appearance in two-dimensionalcross sections of reflectivity fields (Mapes and Houze1993), the air motions in the intermediary and strati-form stages have qualitatively similar divergence pro-files, characterized by a mid- to upper-level layer of

convergence sandwiched between upper and lowerdivergence layers. This three-layered structure differsfundamentally from the two-layered structure seen inthe convective region (Fig. 12a). The three divergenceprofiles sketched in Figs. 12a–c are entirely consistentwith the three observed profiles in Figs. 11a–c. Fromthis we infer that the intermediary and stratiform pro-files both contribute to the slow mode of atmosphericresponse identified by Mapes and Houze (1995, Fig.9), while the convective region profiles (Fig. 11a) con-tribute to the fast mode. To put it another way, the first-order dynamical distinction to be made in evaluatinga tropical convective precipitation system is betweenactively convective portions (like Fig. 11a) and allother intermediary and stratiform portions, which con-tain old convective elements and well-developedstratiform rain.

9. Implications for identifyingconvective and stratiformprecipitation

Because precipitation is a measure of the latent heatof condensation released into the air, and since heat isthe result of upward air motion, the precipitation fieldis diagnostic of the rearrangement of the mass field bytropical convection. This relationship of the heatingand mass fields to the precipitation field is a primarymotivation of the Tropical Rainfall Measuring Mis-sion (TRMM; Simpson 1988). It is now evident thatthe large-scale mass field responds primarily to thetwo-layered mass divergence profile in convectiveprecipitation regions and the three-layered profile instratiform and intermediary precipitation regions. Touse the precipitation field diagnostically, it is neces-sary to separate the convective component of the pre-cipitation from the remainder. This exercise, whileoften referred to as “convective/stratiform separation,”is simply to separate the convection portion of the pre-cipitation from the remainder of the precipitation. Forthis reason, convective/stratiform separation methodsthat seek the stratiform component first and assign theremainder of the precipitation to the convective cat-egory are risky since they are likely to combine theintermediary precipitation (with its three-layered massdivergence profile) with the convective (two-layeredmass divergence profile) to produce a kinematicallynonuniform category of rainfall.

Figure 13 illustrates a method (Churchill and Houze1984; Steiner et al. 1995; Yuter and Houze 1997a,b)


that seeks the convective precipitation first and assignsthe remainder to the “stratiform” category (whichshould really be called stratiform plus intermediary).Each pixel of a Cartesian reflectivity field is examinedand declared a convective center if it exceeds a pre-scribed threshold of reflectivity or if it exceeds thereflectivity in a region surrounding the pixel (~ 10 kmin diameter) by a specified factor, which increases asthe reflectivity at the convective center decreases. If apixel is declared a convective center, then a 4–5-kmdiameter region surrounding the convective center isdefined as convective precipitation. There is no at-tempt to define a convective region smaller than 4–5km in diameter. Often several convective centers areclose together and the convective regions overlap, andthe union of these convective areas may define a con-tiguous region of convective precipitation tens to evenhundreds of kilometers across. Thus, the separationmethod does not attempt to identify individual up- anddowndraft regions, which may be smaller in scale than4–5 km, but rather attempts to delineate general re-gions in which active convection, like that idealizedin Fig. 1a, predominates and hence has a mass diver-gence profile like that in Fig. 12a.

The method illustrated by Fig. 13 is very practicalbecause it is applied to the reflectivity field alone. Itdoes not assume any direct knowledge of the in-cloudair motions. The basic premise of the method is thatif the reflectivity field is either very intense or pock-marked with cores of intense reflectivity, then the airmotions must be more like those in Fig. 1a than thosein Fig. 1d. Steiner et al. (1995) and T.-C. Chen (1996,personal communication) have tested this convective/stratiform separation method on datasets for whichhigh-resolution dual-Doppler radar measurements pro-vided simultaneous measurements of the vertical ve-locity and radar reflectivity so that the vertical veloc-ity data could be used as an independent test of theconvective/stratiform separation. They first deter-mined the convective and stratiform regions by apply-ing the method to the reflectivity field. Then they ex-amined CFADs of the vertical velocity in these re-gions. The statistics of the vertical velocity showedthat the air motions in the regions designated convec-tive had a wide distribution of vertical velocities, withpeaks > ~ 10 m s−1. In the regions that the algorithmdesignated stratiform, the distribution was narrow,with practically no vertical velocities exceeding2 m s−1 in absolute value, with most of the values< 1 m s−1. These statistics confirm that the air motionsin the regions identified as convective by the reflec-

tivity algorithm were like those of Fig. 1a, while theair motions in the stratiform regions were like thoseof Fig. 1d. More tests of this type are needed.

10. Conclusions

In extratropical cyclones, lifting of air is generallywidespread and weak, the precipitation produces lay-ered radar echoes, and the precipitation is traditionallycalled “stratiform.”5 This conceptual model contrastswith that of convective “cells,” which are verticallyoriented cores of high reflectivity, embedded in cumu-lonimbus, which form as a product of free convectiondriven by buoyancy. When radar observations in theTropics in the early 1970s showed stratiform radarechoes alongside convective cells, a paradox arose:stratiform precipitation, thought previously to be aproduct of widespread stable lifting, was falling fromtropical cumulonimbus clouds as a product of buoy-ant convection. Stratiform precipitation and convec-tion could no longer be viewed as mutually exclusive.

This paper suggests resolving this paradox by us-ing the noun convection to describe the overturning ofthe atmosphere that is required to neutralize the verti-cal distribution of moist static energy—just as convec-tion in an incompressible fluid neutralizes an unstablevertical distribution of density—and by using the ad-jectives convective and stratiform to describe the dif-ferent types of precipitation seen within a given regionof convection-generated cumulonimbus.

Within a region of precipitation where the convec-tion is young and vigorous, strong updrafts producecells of heavy rain, seen on radar as locally vertically

5There are several exceptions to the generally stratiform natureof frontal precipitation. The stratiform precipitation often containsweak embedded convective cells, often confined to a potentiallyunstable layer aloft. Release of conditional symmetric instabilitywithin the stratiform cloud mass may lead to embedded lines ofsomewhat heavier precipitation. Embedded cells aloft and linesof release of conditional symmetric instability do not change thefundamentally stratiform nature of the precipitation but, rather,give it a texture. Sometimes, however, the frontal cloud systemcontains a “narrow cold frontal rainband,” in which strong, con-vective-like air motions are forced by the convergence along thesurface wind-shift line. The narrow cold-frontal band is usuallyembedded in a much broader region of predominantly stratiformclouds and precipitation. Sometimes the frontal convergence trig-gers a squall line, in which a deep layer of instability in the warmair is released by the frontal lifting and leads to convective airmotions that dominate over and replace the frontal air motions.For further details, see chapter 11 of Houze (1993).

2194 Vol. 78, No. 10, October 1997

oriented cells of high reflectivity (Figs. 1a–c). Withina region of precipitation where the convection is olderand much less vigorous, a residual population ofweakened updrafts and downdrafts produces precipi-tation that appears stratiform on radar (Figs. 1d–f). Asvigorous cells weaken, they pass through an interme-diary stage of development, when two-dimensionalcross sections of radar data are ambiguous but statis-tics of the reflectivity and vertical air velocity in athree-dimensional volume of space indicate that theradar echoes are basically stratiform; that is, they havea narrow distribution of reflectivity with a vertical dis-tribution suggestive of precipitation particles grow-ing mainly by vapor deposition and vertical veloci-ties < ~ 1 m s−1. In contrast, the statistics of the ech-oes in a three-dimensional volume containing young,vigorous convection show a broad distribution ofreflectivity. Particles there appear to be able to growby collection (coalescence and riming), which ac-counts for the existence of high reflectivities in theoutliers of the reflectivity distribution. Statistics of the

air motions confirm that while most of the vertical ve-locities in the convective regions are relatively small,there are a few large vertical velocities (> ~ 2 m s−1)capable of carrying particles upward and generatinghigh cloud liquid water contents for the growing par-ticles to collect.

Large-scale dynamics of the tropical atmosphere areprofoundly affected by cumulonimbus clouds. How-ever, the effects of the clouds upon the atmosphere aredistinctly subdivided such that part of the atmosphericresponse derives from the young vigorous convectiveprecipitation areas, characterized by radar echoes in theform of cells, while another part derives from the olderweaker part of the convection, where the precipitationechoes are intermediary and stratiform. The young,vigorous cells induce an atmospheric response that istwo layered, with air converging into the active con-vection at low levels and diverging aloft. The older,weaker intermediary and stratiform precipitation ar-eas induce a three-layered response, in which the en-vironment converges into the weak precipitation areaat midlevels and diverges out from it at lower and up-per levels.

These results suggest that it is important to sepa-rate the convective precipitation from all the rest fortwo reasons:

1) The microphysical growth processes at work in theconvective precipitation areas are profoundly dif-ferent from those in the intermediary and stratiformprecipitation areas. In convective precipitation ar-eas, precipitation particles increase their mass pri-marily by collection of cloud water (coalescenceand/or riming). In intermediary and stratiform pre-cipitation areas, the precipitation particles increasetheir mass by vapor diffusion. If high-resolutionnumerical prediction models are to forecast accu-rately the location, amount, and type of precipita-tion, they must represent (either parametrically orexplicitly) the microphysical growth processesaccurately. Radar data can validate these modelsonly if the echoes are accurately subdivided toseparate the active vigorous convective echoesfrom the remainder of the precipitation.

2) Global climate and circulation models must simu-late the role of convection accurately. To validatesuch models, satellite observational programs suchas TRMM (Simpson 1988) are attempting to mapthe precipitation over the globe three dimension-ally. Instrumentation in TRMM will determine thevertical as well as the horizontal structure of pre-

FIG. 13. Schematic illustrating a method for separating theconvective component of precipitation from the rest of theprecipitation. The illustration assumes a 2-km Cartesian gridspacing for the data points; however, the method may be appliedat different resolutions and could be adapted to polar coordinates.The algorithm examines the reflectivity at each grid point. If thatreflectivity exceeds some prescribed high value, or if it exceedsthe mean reflectivity in a “background region” (lightly shadedregion, ~ 10 km in diameter) surrounding the point by somespecified amount, the point is declared a convective center. If thepoint is declared a convective center, then that point and all thepoints in a small region (~ 4–5 km in diameter) surrounding theconvective center constitute a “convective region” centered on thepoint. This exercise is repeated for each data point in the grid andthe union of all the convection regions makes up the totalconvective precipitation region. [Adapted from Steiner et al.(1995).]


cipitation throughout the Tropics. The primary goalis to use this information to determine the four-di-mensional (i.e., temporal and spatial) patterns oflatent heating over the whole Tropics. Algorithmsapplied to the TRMM data will attempt to identifythe two principal modes of heating, by separatingthe convective component of the precipitation, andhence the associated heating, from the remainder.The success of TRMM will hinge on the accuracyof this separation.

From 1) and 2) above, it is clear that the distinctionbetween young convective precipitation and olderstratiform precipitation in convection-generatedclouds is important in order to forecast precipitationaccurately and to evaluate the effects of tropical con-vection on the global circulation. These applicationsare of great practical significance. The processes rep-resented in Fig. 1 remain, however, without muchquantitative documentation. If models for numericalprediction and climate simulation are going to im-prove, these processes must be measured and quanti-fied. The only viable way to do this is to mount fieldprojects that can determine the dynamical and micro-physical processes on space and time resolutions ad-equate to resolve individual convective cells and dis-tinguish unambiguously the major cloud microphysi-cal mechanisms, especially at all altitudes above the0°C level. For these projects, instrumentation and ob-servational platforms must be developed to improveon our current capabilities to measure vertical airmotion, water vapor, temperature, pressure, and iceand liquid particle mass, shapes, and concentrations.Without such empirical improvements in our knowl-edge base, it is hard to foresee real progress in the abil-ity to model processes in the atmosphere that are sen-sitive to convective and stratiform precipitation pro-cesses in cumulonimbus clouds.

Acknowledgments. This paper is based on the author’s presen-tation to the Tropical Rainfall Measuring Mission U.S. ScienceTeam Meeting, Greenbelt, Maryland, July 1996. The author ben-efited greatly from the comments of Dr. Pauline Austin, Profes-sor Colleen Leary, Dr. Brian Mapes, Professor Bradley Smull, andDr. Sandra Yuter. Candace Gudmundson edited the manuscriptand Kay Dewar drafted some of the figures. This research has beensupported by NOAA Cooperative Agreement NA67RJ0155(JISAO Contribution 351), National Science Foundation GrantATM-9409988, and NASA Award NAGS-1599.

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