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From William R. Cotton, George H. Bryan, Susan C. van den Heever, TheParameterization or Modeling of Microphysical Processes in Clouds. In: Storm and

Cloud Dynamics, Vol 99, International Geophysics Series, Renata Dmowska,Dennis Hartmann, H. Thomas Rossby. USA: Academic Press; 2011, pp. 87–142.

ISBN: 978-0-12-088542-8c© Copyright 2011 Elsevier Inc.

Academic Press

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Author’s personal copy

Chapter 4

The Parameterization orModeling of MicrophysicalProcesses in Clouds

4.1. INTRODUCTION

As we have seen, the formulation of a model of a cloud or field ofclouds requires a number of value judgments or compromises. The need forcompromise, however, becomes most obvious when one is faced with the taskof formulating models of the microstructure of clouds. If a modeler with accessto the most advanced levels of computer power is developing a 3-D cloudmodel, the modeler is likely to come to the conclusion that a sophisticated,explicit prediction of the evolution of cloud microstructure is either impossibleor impractical.

The alternative is to develop a simple parameterization of cloudmicrophysical processes. As a general approach, detailed theoretical modelsand/or experimental data are used to formulate parameterizations of thephysics. Ideally, the parameterizations should capture the essence of the knownmicrophysics in simple formulations of the processes. The problem is that, insome cases, the physics is not sufficiently well enough known, or is too complex,to fully capture its essence in simple formulations.

In this chapter we review briefly the concepts and general theory of themicrophysics of clouds. We then summarize approaches to parameterizing cloudmicrophysical processes. We conclude by discussing the interaction of cloudmicrophysical processes and the dynamics of clouds.

4.2. GENERAL THEORY OF THE MICROPHYSICS OF CLOUDS

By “warm” clouds we refer to clouds in which the ice phase does not play asignificant role in either the thermodynamics or precipitation processes. In gen-eral, the term refers to clouds whose tops are no colder than 0 ◦C. However, thephysical processes that are prevalent at temperatures warmer than 0 ◦C can alsooperate quite effectively at colder temperatures or in supercooled clouds. Thus,we will not limit our discussion to clouds that are entirely warmer than 0 ◦C.

87

Author’s personal copy88 PART I Fundamental Concepts and Parameterizations

Noting the differences in droplet concentration in cumuli formed in maritimeand continental air masses, Squires (1956, 1958) introduced the concept ofcolloidal stability of warm clouds. He pointed out that similar clouds formingin a maritime air mass are more likely to produce rain than clouds forming ina continental air mass. Thus, maritime clouds are less colloidally stable thanare their continental counterparts. The relationship between the cloud dropletconcentration and the cloud nucleus population was demonstrated by Twomeyand Squires (1959) and Twomey and Warner (1967). Hence, in a nucleus-rich continental air mass, a given liquid-water content must be distributedover numerous small droplets having small collection kernels or collectioncross sections (i.e. low terminal velocities, collection efficiencies, and cross-sectional areas). Thus, in general, the collision and coalescence process isinhibited in nucleus-rich continental air masses. These concepts form the basisof many of the parameterizations of warm-cloud processes that we shall discussshortly. The fundamental premise of many of the parameterizations is thatthe cloud droplet concentration or activated cloud condensation nuclei (CCN)concentration, at cloud base, determine whether or not a cloud will precipitate.

Johnson (1980) pointed out that the cloud-base temperature also influencesthe activation of cloud droplets. Other things being the same (i.e. aerosoldistribution, cloud-base updraft velocity), clouds with colder cloud baseswill activate more cloud droplets than those having warmer bases. This isa consequence of the nonlinear variation of saturation vapor pressure withtemperature, which results in higher peak supersaturations in cold based cloudsthan in warm-based clouds that are otherwise the same. It is the direction of thiseffect that is most interesting, because the temperature effect will accentuatethe tendency for colloidal stability in continental clouds if those clouds alsohave cold bases (a common occurrence in mid-latitude, continental regions).The aerosol distribution and updraft velocity at cloud base, however, remain asthe most important factors controlling the concentration of activated dropletsand, hence, the colloidal stability of a cloud.

Cloud-base temperature also figures in a cloud’s colloidal stability becausea cloud with a warm cloud-base temperature has a larger saturation mixingratio at cloud base. Other things being the same (i.e. cloud depth, activatedCCN concentration, etc.), a warm-based cloud will have a greater potentialfor producing a significant amount of condensed liquid-water and, therefore,a greater chance of generating a few big droplets.

For some time it was thought that collision and coalescence could notproceed until the radius of droplets exceeded 19 µm (Hocking, 1959). Latercalculations of collision efficiency by Klett and Davis (1973), Hocking andJonas (1970), and Jonas and Goldsmith (1972) suggest that droplets smallerthan 19 µm in radius do exhibit finite collection efficiencies. Due to their smallfall velocities and cross-sectional areas, however, the rate of collision amongdroplets of such a small size is very low. Thus, it is still thought that a few largerdroplets r > 20 µm must form in a cloud in order to initiate significant growth

Author’s personal copyChapter 4 Modeling Cloud Microphysics 89

rates through the relatively random collisions among small, comparably sizeddroplets. This initial phase of collision and coalescence has been modeled as astochastic process (Telford, 1955; Gillespie, 1972) or what is now referred to asquasistochastic process (Berry, 1967).

Since condensation theory for a smooth, unmixed updraft predicts anarrowing of the droplet spectrum with time (Howell, 1949; Mordy, 1959;Neiburger and Chien, 1960; Fitzgerald, 1974), the search continues to explainthe formation of larger droplets that can sustain vigorous collision andcoalescence growth of precipitation.

There are four hypothesized mechanisms to explain the formation of dropletslarge enough to be considered precipitation embryos:

• Turbulence influences on condensation growth.• Role of giant cloud condensation nuclei.• Turbulence influence on droplet collision and coalescence.• Radiative cooling of droplets to form precipitation embryos.

4.2.1. Turbulence Influences on Condensation Growth

Simple adiabatic parcel models assume a uniform updraft and develop aconstant droplet concentration a few meters above cloud base. Convectiveclouds are, in fact, made up of a series of updrafts of diverse intensities.The least vigorous updraft at cloud base also produces the lowest dropletconcentration during activation of the available CCN. Turbulence furthercontributes to entrainment of dry environmental air in clouds, hence reducingLWC below adiabatic values and diluting the droplet concentration belowits initial value after CCN activation. Convective cells with a significantlyreduced concentration might generate bigger droplets than adiabatic cores, ifthey are further experiencing a convective ascent (Baker et al., 1980; Telford andChai, 1980; Telford et al., 1984). Turbulence, however, also contributes to thecontinuous mixing of the convective cells, hence smoothing out their differencesin terms of concentration and droplet growth.

Airborne measurements performed in cumulus clouds with a high resolutiondroplet spectrometer Fast-FSSP reveal the occurrence of very narrow dropletspectra in cloud cores (Brenguier and Chaumat, 2001). However, droplet spectraare often much broader than the narrow adiabatic reference, with droplet sizesextending from zero to the maximum predicted by the adiabatic model. Thisreflects the impact of mixing between convective cells that have experiencedvarious levels of dilution with the entrained air. There are also airborneobservations in stratocumulus clouds showing negative correlations betweendroplet concentration and droplet sizes (see for example Fig. 7 in (Pawlowskaand Brenguier, 2000). This corroborates the hypothesis that fluctuations ofthe updraft intensity at cloud base or the ascent of diluted cloud cells mightcontribute to the formation of droplets bigger than the adiabatic prediction basedon the mean droplet concentration. In deeper clouds, airborne cloud traverses at


a given level all look rather similar in term of droplet concentration and sizes,suggesting that concentration/size correlations progressively dissipate due tocontinuous mixing.

Numerical simulations (Vaillancourt et al., 2002; Shaw et al., 1998) suggestthat turbulence may also generate concentration fluctuations at the microscaleby inertia in the regions of high vorticity, hence leading to superadiabatic dropletgrowth in the microcells with the lowest concentrations. This hypothesis has notbe supported by in situ measurements of the droplet spatial distribution at themicroscale (Chaumat and Brenguier, 2001).

In summary, the observational evidence that turbulence contributes tothe formation of precipitation embryos by enhanced condensation growth islacking. In fact, stochastic processes induced by turbulence are not likelyto enhance droplet growth because condensation is a cumulative process. Toexperience superadiabatic growth, droplets need to remain isolated in regionsof higher supersaturation for a significant part of their ascent. The odds of thisare low as mixing continuously redistributes droplets in the cloud. Turbulence ismore likely to affect a discontinuous process like collision, since once dropletscoalesce, they cannot be separated.

4.2.2. Role of Giant Cloud Condensation Nuclei (GCCN)

Observations reported by Woodcock (1953), Nelson and Gokhale (1968),Hindman (1975), Johnson (1976, 1982), and Hobbs et al. (1980) have shown thepresence of potentially significant concentrations of aerosol particles of sizesas large as 100 µm. Their concentrations are about 10−3 cm−3 (Woodcock,1953), that is, about one in 105 or 106 CCN are giant particles. Nevertheless,these particles can have a significant effect on the development of precipitationby serving as coalescence embryos (Johnson, 1982; Feingold et al., 1999; Yinet al., 2000). The droplet that forms is large enough for coalescence to startimmediately even before the droplet reaches its critical size based on the Kohlerequation. This can occur if the nuclei are completely soluble (e.g. sea-saltparticles) or are mixed particles with a soluble coating (e.g. mineral dust witha coating of sulfate, (Levin et al., 1996) or are very large and wettable. Thepresence of GCCN on precipitation formation has been investigated in a numberof cloud resolving models, which show that their contribution to rain formationmay be appreciable in polluted clouds but has little influence in clouds formingin clean air masses (i.e. Feingold et al., 1999; Khain et al., 2000).

4.2.3. Turbulence Influence on Droplet Collision andCoalescence

Turbulence can influence the collision and coalescence process in three ways:

• By enhancing collision efficiencies.• By enhancing the collection kernels.


• By producing inhomogeneities in droplet concentration.

Collision efficiencies are generally calculated in laminar or stagnant flow.In turbulent flow droplets will be accelerating and thereby be able to crossstreamlines more efficiently than in laminar flow resulting in enhanced collisionefficiencies. Large droplets, having more inertia, will be affected more byturbulence than smaller drops. Calculations by Koziol and Leighton (1996)suggest that this effect is small for droplets smaller than 20 µm diameter.However, turbulence can also cause fluctuations on vertical fall speeds andhorizontal motions, such that the collection kernel is enhanced (Pinsky andKhain, 1997; Khain and Pinsky, 1997), relative to that defined in laminarflow. Because the collection rate is proportional to the square of dropletconcentrations, inhomogeneities in droplet concentrations due to turbulence canproduce enhanced regions of collection where the droplet concentrations arelocally enhanced in, say, regions of low vorticity (Pinsky and Khain, 1997).

The problem is that there is little known about the details of turbulencein real clouds on the scales of a few centimeters and less. We know thatinhomogeneities exist on those scales (Baker, 1992), but when it comes to eitherlaboratory simulations or theoretical calculations, our ability to simulate highReynolds number turbulence and its effects on droplet collection is still veryrudimentary.

4.2.4. Radiative Cooling of Droplets to Form PrecipitationEmbryos

Consider a population of droplets that resides near a cloud top for a sufficientlylong time. Those droplets will emit radiation to space quite effectively if theatmosphere above is relatively dry and cloud free. As a result, the droplets willbe cooler than they would be without considering radiative effects. This meansthat the saturation vapor pressure at the surface of the droplet will be loweredand the droplets will grow faster.

But radiation cooling is proportional to the cross sectional area of a dropletso that its effect is much greater on larger droplets than small ones (Roach, 1976;Barkstrom, 1978; Guzzi and Rizzi, 1980; Austin et al., 1995). In fact, Harringtonet al. (2000) have shown that in a marine stratocumulus environment, whendroplets are competing for a limited supply of water vapor, the larger dropletsgrow so rapidly by radiative enhancement that droplets smaller than 10 µmin radius evaporate producing a bimodal size spectrum. This process is onlyeffective in clouds where droplets reside near the cloud top for time scales of12 minutes or longer such as fogs, stratus, and stratocumulus. Cumulus cloudswith vigorous overturning expose droplets to space for too short a time.

Whatever the nature of the process of formation of embryonic droplets, theportions of a given cloud most favorable for the initiation of precipitation inwarm clouds are the regions of highest liquid-water content. Twomey (1976)showed that if locally enhanced regions of LWC comprise only 1% of the


cloud volume, and exist for periods of a few minutes, such regions can producesignificant concentrations of large drops averaged over the entire volume of thecloud. Thus, the presence of protected updrafts having nearly wet adiabaticliquid-water contents (Heymsfield et al., 1978) can have significant bearingupon the initiation of precipitation in warm clouds. Of course, the ultimateamount of rainfall from a given cloud is controlled by the overall time-spacecharacter of its updrafts and its liquid-water content.

Langmuir (1948) suggested that once raindrops grow to a critical size ofapproximately 6 mm in diameter, they will break up due to hydrodynamicinstability. He hypothesized that each breakup fragment will act as a newprecipitation embryo which can grow to breakup size and create more raindropembryos. He referred to this process as the chain reaction theory of warm-rain formation. Other observations (Blanchard, 1948; Magarvey and Geldhart,1962; Cotton and Gokhale, 1967; Brazier-Smith et al., 1972; McTaggert-Cowanand List, 1975) have suggested that collisions among droplets of the order of2–3 mm in diameter and smaller can initiate breakup.

Computations of the evolution of raindrop spectra reported by Brazier-Smith et al. (1973), Young (1975), and Gillespie and List (1976) have indicatedthe greater importance of collision-induced breakup over spontaneous breakupto the evolution of raindrop spectra. Srivastava (1978) calculated that forrainwater contents M > 1 g m−3, collision breakup results in raindrop sizedistributions which are approximately constant in slope and have an interceptof the distribution function which is proportional to M . Using a numericalcloud model, Farley and Chen (1975) have concluded that a necessary conditionfor the development of a Langmuir chain reaction requires that a cloud mustdevelop sustained updrafts in excess of 10 m s−1

4.3. APPROACHES TO MODELING CLOUD MICROPHYSICS INCLOUD MODELS

4.3.1. Bin-resolving Microphysics

The approach to modeling microphysics in clouds ranges from explicit bin-resolving approaches in which the evolution of droplet and ice particle spectraare explicitly resolved, to bulk approaches in which the size distribution ofhydrometeors is determined by a specified basis function in which one or moremoments of the basis function is predicted.

The most simple class of model in which droplet spectra are explicitlyresolved is the Lagrangian (or moving mass-grid) method. In this methodparticles at discrete sizes and concentrations follow the growth by condensationon a moving mass grid. This approach eliminates numerical diffusion and allowsfor a smooth transition from aerosol to haze to cloud droplets without artificialdistinctions between these classes. The cloud supersaturation is calculated basedon source (cooling, which is related to updraft velocity) and sink (condensation)terms, enabling accurate determination of the number of activated droplets.


These models typically focus on the initial growth phases from haze to dropletand include detailed representation of aerosol sizes and compositions (Mordy,1959; Fitzgerald, 1974; Facchini et al., 1999; Feingold and Kreidenweis, 2000;Feingold and Chuang, 2002; Lohmann, 2004). They frequently also consider theeffect of trace gases on activation (Kulmala et al., 1993). Aqueous productionof sulfate has also been represented in studies that examine the effects ofcloud processing on the aerosol size distribution (Hegg et al., 1991; Bower andChoularton, 1993; Feingold and Kreidenweis, 2000). The Lagrangian methodis used almost exclusively in kinematic cloud parcel models, where a parcel ofair is moved either adiabatically or according to some known trajectory througha cloud. It is not easily adapted to the study of growth by processes such ascollision-coalescence, and it is not suitable for general application in Euleriandynamical models.

Another class of models that explicitly resolve the evolution of hydrometeorsize spectra is the so called bin-resolving technique. Assuming that the dropletspectrum is continuous on the scale of the averaging domain of a cloud model,one can formulate integral-differential equations that describe the evolution ofdroplet spectra. For example, we can formulate a prognostic equation for thevariation of the spectral density f (x) of cloud droplets of mass x to x ± δx/2at a given geometric position and at a given instant. An integral differentialequation describing the evolution of the droplet spectrum takes the form

∂ f (x)

∂t= N (x)−

∂ [x f (x)]∂x

+ G(x)|gain + G(x)|loss

+ B(x)|gain + B(x)|loss + τ(x), (4.1)

where N represents nucleation, G represents collection, B represents breakup,and τ represents the sum of both mean and turbulent transport processes.

The first term on the right-hand side of Eq. (4.1) is the production of dropletsof mass x by the nucleation of such droplets on activated CCN. This termappears in Eq. (4.1) only if the droplet spectrum f (x) is truncated at some smalldroplet mass.

The second term on the right-hand side of Eq. (4.1) is the divergence off (x) due to continuous vapor mass deposition on droplets growing at the ratex , where x is a function of the droplet mass, its solubility in water, and thelocal cloud supersaturation as well as other factors (Byers, 1965; Mason, 1971;Pruppacher and Klett, 1978). If one chooses to extend the droplet spectrum toinclude soluble particles of mass xs , the nucleation term (N) would be includedexplicitly in the second term. However, since x is a function of the solubility ofa droplet, one should use a two-dimensional density function f (x, xs), as wasdone by Clark (1973).

The third and fourth terms on the right-hand side of Eq. (4.1) represent,respectively, the gain and loss integrals of f (x) due to the collision andcoalescence of cloud droplets. The gain and loss terms were formulated by


Berry (1967) as follows

G(x)|gain =12

∫ x

0K (xc, x ′) f (xc) f (x ′)dx ′ (4.2)

where xc = x − x ′ and

G(x)loss = −

∫∞

0K (x, x ′) f (x) f (x ′)dx ′. (4.3)

The term K (x, x ′) is the collision cross section, or collection kernel, oftentaken to be

K (x, x ′) = π(Rx + Rx ′)2[v(x)− v(x ′)]E(x, x ′), (4.4)

where E(x, x ′) is the collection efficiency, Rx and Rx ′ are the radii of droplets ofmass x and x ′, and v(x) and v(x ′) are the average terminal velocities of dropletsof mass x and x ′.

Integration of the droplet spectra evolution using Eqs (4.2)–(4.4) has beenreferred to by Gillespie (1975) as the quasistochastic model because it predicts aunique spectrum at a given time and point in space, whereas the pure stochasticmodel predicts fluctuations in the droplet spectrum at a given time and point inspace.

The fifth and sixth terms on the right-hand side of Eq. (4.1) represent,respectively, the gain and loss of spectral density f (x) due to the breakup ofdroplets.

Thus, the breakup of larger droplets whose fragments are of mass x ± δx/2is formulated as

B(x)|gain =

∫∞

xp(x ′) f (x ′)g(x |x ′)dx ′, (4.5)

where p(x ′) is the probability per unit/time that a droplet of mass x ′ to x ′±δx/2will break up due to internal hydrodynamic instability. The function g(x |x ′)represents the number of fragments of mass x ± δx/2 formed by the breakup ofa droplet of mass x ′ ± δx/2.

The loss of f (x) due to the breakup of droplets of mass x ± δx/2 isformulated as

B(x)|loss = −p(x) f (x). (4.6)

It was mentioned earlier that breakup due to collision of raindrops appearsto be the dominant contributor to drop breakup, at least if the rainwater contentsexceed 1 g m−3. Brazier-Smith et al. (1973) formulated collision breakup bycombining the collision and breakup terms in Eq. (4.1) to formulate a general


stochastic interaction equation of the form

G(x)|gain + G(x)|loss + B(x)|gain + B(x)|loss

=12

∫∞

0

∫∞

0K (x |xα, xβ) f (xβ) f (xα)dxαdxβ

− f (x)∫∞

0

∣∣∣∣∫ ∞0

K (xα|x, xβ)xαdxα

∣∣∣∣ f (xβ)

x + xβdxβ , (4.7)

where K (x |xα, xβ) may be considered a generalized interaction kernel relatedto the probability [h(x |xα, xβ)] of forming a droplet of mass x because of theinteraction of droplets of mass xα and xβ . In this instance, the interaction couldrepresent a pure coalescence problem in which case Eq. (4.7) degenerates toEqs (4.2) and (4.3). Alternatively, it could represent a collision event whichpromotes breakup.

A common approach to integrating Eq. (4.1) is to discretize f (x) into 40to 70 elements and then integrate the equations by finite-difference methods(Twomey, 1964, 1966; Bartlett, 1966, 1970; Warshaw, 1967; Berry, 1967;Kovetz and Olund, 1969; Bleck, 1970; Chien and Neiburger, 1972). Great caremust be taken in representing collision-coalescence to avoid numerical diffusionin the mass-transfer equations and rapid (spurious) acceleration of growth toprecipitation-sized particles.

Multi-moment representations of cloud processes have been developed(Tzivion et al., 1987; Hounslow et al., 1988; Chen and Lamb, 1994b) inwhich two or more moments in each individual drop bin are predicted. Thisapproach significantly reduces numerical diffusion and has the added benefit ofconserving more than one moment in each bin of the size distribution.

Another approach to bin-resolved microphysical modeling is the hybridapproach (Cooper et al., 1997; Jacobson, 1999; Pinsky and Khain, 2002)where the advantages of moving grids for condensational growth are combinedwith fixed grid approaches for collection processes. In spite of their largecomputational demands, bin-resolving microphysical models have been appliedto large eddy simulation models (Stevens et al., 1998), mixed-phase threedimensional models of thunderstorms (Fridlind et al., 2004) and even mesoscalemodels (Lynn et al., 2005). However,for complicated cloud or mesoscalemodels, forecast models, and climate models, there still remains a strong desireto develop simplified techniques for predicting the evolution of the dropletspectrum to form rain along with its sedimentation through the cloud.

4.3.2. Bulk Warm-cloud Physics

The concept of bulk microphysics was first introduced by Kessler (1969).He assumed that the raindrop size-spectrum can be represented by a simpleexponential function (with fixed pre-exponent) in which a single-moment ofthe hydrometeor spectra is predicted. Self-collection among cloud droplets is


0.3

0.2

0.1

t = 0

t = 500

t = 0

t = 0

t = 500

t = 1200

t = 700

t = 1400

t = 700

t = 900

t = sec

N0 = 300 cm–3

m = 2.5 g m–3

rγ = 0.25

rγ = 0.25

rγ = 0.25

N0 = 300 cm–3

m = 2.0 g m–3

N0 = 300 cm–3

m = 1.5 g m–3

(a)

(b)

(c)

0.5

0.4

0.3

0.2

0.1

0

0.4

0

g(ln

R)

in g

m–3

/(

ln R

)δ

0.3

0.2

0.1

010 100

Radius (μm)1 1000

FIGURE 4.1 Computed variation in the distribution of water mass density for an initialconcentration of 100 cm−3, a radius dispersion of 0.25, and LWCs of (a) 2.0, (b) 1.5, and (c)1.0 g m−3. (From Cotton (1972a))

parameterized using an autoconversion formulation, and large hydrometeorssuch as raindrops are assumed to collect smaller drops by continuous accretion.Moreover, all classes of large hydrometeors are assumed to fall with a constantfall speed, usually a water content-weighted fall speed. To illustrate thisapproach consider Fig. 4.1 in which condensed liquid-water is partitioned intotwo domains. Note that the area under the curve is the liquid-water content.

Generally, we distinguish between cloud droplets, assumed to havesufficiently small terminal velocities to generally move with air parcels, andraindrops, which have significant settling speeds v(R), where R is the dropradius. Kessler’s (1969) identification of an arbitrary threshold, separatingcloud droplets falling at small terminal velocities from higher terminal velocityraindrops, is illustrated as a vertical dashed line in Fig. 4.3. To the left of thevertical dashed line, water vapor is condensed on low-terminal-velocity clouddroplets, where droplet collection forms raindrops which are on the right-hand


side of the line. Berry and Reinhardt (1974) demonstrated that a natural breakbetween cloud and raindrops occurs at a radius of 50 µm.

4.3.3. Conversion Parameterizations

Kessler (1969) first developed a simple parameterization of the rate ofautoconversion of condensed liquid water from cloud droplets, having liquid-water content m (mass/volume), to raindrops, having water content M . Heassumed that the rate of conversion was a linear function of m for liquid-watercontents greater than 1.0 g m−3.

Transformed into mixing ratio quantities, Kessler’s autoconversion formulacan be written

CNcr = −ρ−1o (dm/dt) = K1(rc − a), (4.8)

where K1 > 0 if rc > a and K1 = 0 if rc ≤ a. In our notation we haveadopted the subscript cr , which indicates that raindrops having mixing ratios rrare supplied by cloud droplets having mixing ratios rc.

Over the years there has been a variety of autoconversion formulationsthat have either been ad hoc [e.g. Kessler (1969), Manton and Cotton (1977),Cotton et al. (1986)], or derived from detailed bin-microphysics simulations[e.g. Berry (1967), Cotton (1972a,b), Berry and Reinhardt (1974), Beheng(1994), Khairoutdinov and Kogan (2000)]. Liu and Daum (2004) showed thatmany of the Kessler type of autoconversion formulations are special cases ofa more general formulation. Note that for a given liquid-water content theseformulations differ by several orders of magnitude.

Application of the conversion rate formulas to deep convective cloudsreveals that these formulas serve mainly as triggers for warm-rain initiation.Once precipitation-size droplets form, the rain formation process is dominatedby raindrops accreting smaller cloud droplets. As such, the absolute magnitudeof the conversion rate may not be very important.

Only in stratocumuli and shorter lived, low-liquid-water cumuli will themagnitude of the conversion rate be essential to determining if such a cloudwill produce precipitation.

4.3.4. Parameterization of Accretion

Once embryonic precipitation particles are formed, Kessler (1969) hypothesizedthat the water content converted to rain is distributed in an exponentialdistribution function formulated by Marshall and Palmer (1948) as

N (D) = N0e−λD, (4.9)

where N (D) represents the number of raindrops per unit/volume of diameterD ± δD/2. Figure 4.2 illustrates the Marshall-Palmer distribution function.


N0

ln N

(D)/

Dln

(R

)/ R

φ

δδ

3

2λ

λ

λ λ

1

D

(a)

(b)

= = CONSTANT

R

N03N02N01

lRm

FIGURE 4.2 Schematic illustration of the exponential drop-size distribution function ofMarshall-Palmer drop-size distribution. (a) Kessler’s model in which N0 is assumed constantand λ varies with rainwater content; (b) Manton-Cotton’s model in which the slope λ = 1/Rmis a constant and N0 = NR/Rm varies with rainwater content.

Kessler assumed that the rate of mass growth of raindrops is primarily byaccretion of cloud droplets of the form

(d/dt)[x(D)] = (πD2/4)EVDρ0rc, (4.10)

where x(D) is the mass of a raindrop of diameter D, VD is the terminal velocityof a raindrop of diameter D, and E represents an average collection efficiencybetween raindrops and cloud droplets.

The rate of change of rainwater mixing ratio by accretion or collection ofcloud droplets is then

CLcr =1ρ0

∫∞

0

ddt[x(D)]N (D)dD, (4.11)

or

CLcr =1ρ0

∫∞

0

πD2

4EVDρ0rc N (D)dD. (4.12)


Kessler further assumed that V D is given by

V D = 130.0D1/2 m s−1, (4.13)

from Spilhaus (1948). Then, after substituting Eq. (4.19) into Eq. (4.18) andintegrating, we find that

CLcr =130.0ρ

4N0 Eρ0rc

0(3.5)

λ3.5 , (4.14)

Under the assumptionthat N0 is a constant, the total rainwater mixing ratiorr may be obtained as

rr =1ρ0

∫∞

0X (D)N (D)d(D). (4.15)

After substituting Eq. (4.15) and recognizing that

X (D) = πD3ρ1/6, (4.16)

Eq. (4.21) becomes, after integrating,

rr = πρ1 N00(4)ρ06λ4. (4.17)

Solving Eq. (4.23) for λ and substituting into Eq. (4.20) gives us

CLcr =

{130.0π0.125

4

[6

0(4)

]0.875

0(3.5)

}× N 0.125

0 ρ1.8750 Ercr0.875

r . (4.18)

Equation (4.24) shows that raindrops accreting cloud droplets produce arainwater mixing ratio at a rate proportional to rcr0.875

r .Over the years there has been a number of refinements and modifications

to the simple Kessler formula. They include assuming that the slope of theexponential function is a constant (Manton and Cotton (1977), see also Tripoliand Cotton (1980)) to the use of gamma or log-normal basis functions Clark(1976); Clark and Hall (1983); Nickerson et al. (1986); Ferrier (1994); Meyerset al. (1997); Reisner et al. (1998); Seifert et al. (2006); Milbrandt and Yau(2005a,b), and then explicitly predicting the evolution of two or more moments.The advantage of the multimoment schemes is that they predict numberconcentration, mass mixing ratio (and sometimes higher order moments) andtherefore are able to derive the broad features of the drop size distribution. Inso doing they improve the representation of growth processes and precipitationformation.


Small cloud droplet distributionMean diameter = 20 micronsNumber concentration = 100 cm–3

Large cloud droplet distribution (x 50)Mean diameter = 60 micronsNumber concentration = 1 cm–3

Rain droplet distribution (x 10000)Mean diameter = 0.3 cmNumber concentration = 10–4 cm–3B

inne

d nu

mbe

r co

ncen

trat

ion

(cm

–3)

1.0

0.8

0.6

0.4

0.2

0.00.001 0.01 0.1

Log-normal droplet diameter (cm)

1

FIGURE 4.3 Droplet spectra for the rain and dual cloud droplet hydrometeor categories.For comparison to the small cloud droplet mode, the number concentration for the given diameteris exaggerated for the rain and large cloud droplet modes. (From Saleeby and Cotton (2004))

Departing from the Kessler bulk parameterization philosophy, the bulkscheme in RAMS (Regional Atmospheric Modeling System, Cotton et al.2003) essentially emulates a full-bin microphysics model. The evolution of thismodeling approach can be found in Clark and Hall (1983), Verlinde et al. (1990),Walko et al. (1995), Meyers et al. (1997), Feingold et al. (1998), and Saleebyand Cotton (2004). Instead of using continuous accretion approximations, whichhas been common in cloud parameterizations, Feingold et al. (1998) showed thatfull stochastic collection solutions for self-collection among cloud droplets andfor rain (drizzle) drop collection of cloud droplets can be obtained for realisticcollection kernels by making use of look-up tables. Saleeby and Cotton (2004)refined this approach by introducing two cloud modes, one for newly nucleateddroplets on CCN having diameters less than 40 µm, and a second larger dropletmode from 40 to 80 µm in which giant CCN (GCCN) are nucleated. Theactivation of CCN in RAMS is parameterized using a look-up table derived froman ensemble of a Lagrangian parcel model calculations that considers ambientcloud conditions for the activation of cloud droplets from aerosol.The use oftwo cloud droplet modes provides a quantitative representation of the collectionprocess and permits representing both CCN and GCCN as contributors to thecollection process. Figure 4.3 illustrates schematically the prescribed basisfunctions. This approach has been extended to all hydrometeor class interactionsby collection, including the growth of graupel and hail by riming (Saleeby et al.,2007). The philosophy of bin representation of collection has been extended tocalculations of drop sedimentation (Feingold et al., 1998). Bin sedimentationis simulated by dividing the basis function into discrete bins and then buildinglook-up tables to calculate how much mass and number in a given grid cellfall into each cell beneath a given level in a given time step. This permits the


representation of size-sorting of hydrometeors which is not done in standardbulk schemes.

4.4. FUNDAMENTAL PRINCIPLES OF ICE-PHASEMICROPHYSICS

The representation of ice-phase microphysical processes in a cloud model isgreatly complicated by the variety of forms of the ice phase, as well as bythe numerous physical processes that determine the crystal forms. Moreover, incontrast to the physics of warm clouds, our understanding of ice-phase physicsis far less complete. This means that in many cases the formulation of simpleparameterized models of the ice phase cannot be done using information derivedfrom detailed theoretical/numerical models or from observations.

The physical processes that should be considered in formulating a model orparameterization of the ice phase are as follows:

(1) Primary and secondary nucleation of ice crystals.(2) Vapor deposition growth of ice crystals.(3) Riming growth of ice crystals.(4) Graupel or hail particle initiation from heavily rimed crystals.(5) Graupel or hail particle initiation by the freezing of supercooled raindrops.(6) Graupel or hail particle riming and vapor deposition growth.(7) Graupel or hail particle-particle collision with supercooled raindrops.(8) Shedding of water drops from hailstones growing by wet growth or from

partially wetting ice particles.(9) The initiation of aggregates of ice crystals by collision among ice crystals.

(10) Aggregate collection of ice crystals.(11) Aggregate riming of cloud droplets.(12) Melting of all forms of ice particles.

4.4.1. Nucleation of Ice Crystals

Ice particles can form either homogeneously or heterogeneously on some formof ice nuclei (IN). Homogeneous nucleation can take place either directly fromthe vapor or by freezing of cloud droplets. However, homogeneous nucleationof ice crystals from the vapor, or the chance formation of an embryo of ice-likestructure of critical size, requires very high supersaturations with respect to iceand such low temperatures that it does not take place in the troposphere. Onthe other hand, homogeneous freezing of supercooled droplets by the chanceformation of a cluster of ice-like embryos can occur in the atmosphere.

For homogeneous freezing to occur, enough ice-like water molecules mustcome together within the droplet to form an embryo of ice large enough tosurvive and grow. Because the numbers and sizes of the ice embryos that formby chance increase with decreasing temperature, below a certain temperature(which depends on the volume of water considered) freezing by homogeneous


nucleation becomes a virtual certainty. Homogeneous nucleation occurs in aboutone second at about −41◦C for droplets about 1 µm in diameter, and at about−35 ◦C for drops 100 µm in diameter. An analogous freezing process occurs forunactivated droplets or haze particles at temperatures below −40◦C, a processfor ice formation in cirrus clouds (DeMott, 2002). Hence, in the atmosphere,homogeneous nucleation by freezing generally occurs only in high clouds orhigh latitudes.

The presence of some form of nucleus or mote is required at temperatureswarmer than −35 ◦C, in a process called heterogeneous nucleation. Theprinciple ice nucleation mechanisms are (1) vapor-deposition nucleation, (2)condensation-freezing, (3) immersion-freezing nucleation, and (4) contact-freezing nucleation.

Vapor deposition nucleation refers to the direct transfer of water vapor to anucleus that results in the formation of an ice crystal. Condensation-freezingnucleation refers to the condensation of water vapor on an internally-mixednucleus to form an embryonic droplet, followed by freezing. This is viewedas a two-step process, so the name “condensation freezing” is often used. Inpractice it is not easy to distinguish between these two modes of nucleation.

Immersion freezing refers to the nucleation of a cloud droplet or raindrop onan ice nucleus which is immersed within the drop. Two theories of immersionfreezing have emerged. One theory views the freezing process as stochastic,such that at a given degree of supercooling, not all drops of a population ofdrops will freeze at the same time Bigg (1953a,b, 1955); Carte (1956); Dufuorand Defay (1963). The second theory, called the singular theory, holds thatat a given degree of supercooling, the probability that a drop of a given sizewill freeze depends solely on the likelihood that the drop contains an activefreezing nucleus. This process is independent of time. Laboratory experimentsreported by Vali and Stansbury (1966) have indicated that the freezing process istime dependent, though the amount of time dependence is small. Both theoriespredict that the probability of freezing increases exponentially with the degreeof supercooling and with the volume (or size) of the drop. Thus, for small clouddroplets at small degrees of supercooling, this mechanism of nucleation is notvery effective.

Contact nucleation refers to the nucleation of a supercooled drop by anucleus that makes contact with the surface of the drop in the supercooledstate. Observations have shown that dry particles, such as clays, sand, CuS,and organic compounds which make contact with a supercooled drop, are muchmore effective as contact-freezing nuclei than when immersed within the drop(Rau, 1950; Fletcher, 1962; Levkov, 1971; Gokhale and Spengler, 1972; Pitterand Pruppacher, 1973; Fukuta, 1972a,b). In the atmosphere, contact nucleationcan produce considerable time dependence in the nucleation process, since itdepends both on the probability that an aerosol particle makes contact with asupercooled drop and on the probability that the aerosol particle acts as an activefreezing nucleus. Young (1974a) modeled contact nucleation by naturally and


artificially generated aerosols. He considered cloud droplet scavenging of nucleipanicles by Brownian diiffusion and by the combined effects of thermophoresisand diffusiophoresis.

Brownian diffusion refers to the chance encounter between a supercooledcloud droplet and an active nucleus (aerosol particle) due to the randommotion of both species as a consequence of the thermal bombardment withgas molecules. The rate of collision is a function of the kinetic energyof the air molecules (or temperature) and the mobilities of both species.Thermophoresis refers to a net transport of particles in a thermogradient fromwarm toward colder regions. Diffusiophoresis refers to the net transport ofaerosol particles in the direction of a vapor flux. In the case of a droplet growingby vapor deposition, diffusiophoresis is directed toward the droplet, whereasthermophoresis acts away from the droplet. The reverse is true of an evaporatingdroplet. Young (1974a) has noted that since thermophoresis dominates overdiffusiophoresis, contact nucleation by nuclei in the size range 0.15 < 1.0 µmwould be suppressed in a growing cumulus cloud. He cites observations reportedby Mee and Takeuchi (1968) and Koenig (1962) which indicate that ice is mostprevalent in downdrafts and at the edges of clouds. This gives evidence, hesuggests, that phoretic-contact nucleation is an effective mechanism in naturalclouds.

Historically, ice nuclei concentrations have been assumed as predictors ofice particle concentrations. Laboratory studies in the 1940’s and 1950’s usingexpansion chambers, mixing cloud chambers or aerosol collected on filters,indicated that the variability in ice nuclei concentration was strongly a functionof temperature. Fletcher (1962) derived an empirical relationship relating theincrease of concentration of ice nuclei with decreasing temperature:

NIN = A exp(βTs), (4.19)

where N is the concentration of active ice nuclei per liter of air, Ts is the degreeof supercooling, β varies from about 0.3 to 0.8 and A is about 10−5 liter−1. Forβ = 0.6, Eq. (4.19) predicts that the concentration of ice nuclei increases byabout a factor of 10 for every 4 ◦C decrease in temperature. In urban air, thetotal concentration of aerosol is on the order of 108 liter−1 and only about oneparticle in 108 acts as an ice nucleus at −20 ◦C.

The activity of a particle as a condensation-freezing or a deposition nucleusdepends not only on the temperature but also on the supersaturation of theambient air. Thus, at a given temperature, Gagin (1972), Huffman (1973),and Huffman and Vali (1973) found that the concentration of IN varieswith the supersaturation with respect to ice. The effect of supersaturation onmeasurements of ice nucleus concentrations is shown in Fig. 4.4, where itcan be seen that at a constant temperature the greater the supersaturation withrespect to ice the more particles serve as ice nuclei. The empirical equation to


–20

–20

Ice

conc

(L–

1 )

Ice supersaturation (%)

–20–15

–10

–16–15

–12–10

–7

Cotton et al. (1986)

100

10

1

0.1

0 5 10 15 20 25

FIGURE 4.4 Continuous-flow diffusion-chamber ice-nucleus concentration measurementsversus ice supersaturation from (open square) Rogers (1982), and from (filled square) Al-Naimi and Saunders (1985). Constant-temperature measurement series are indicated and theregression given in Eq. (4.20) is shown. Also presented are constant temperature values predicted bythe deposition-condensation-freezing nucleation model formulated by Cotton et al. (1986). (FromMeyers et al. (1992))

the best-fit line to these measurements is

Ni = exp{a + b[100(Si − 1)]}, (4.20)

where Ni is the concentration of ice nuclei per liter, and Si − 1 is thesupersaturation with respect to ice, a = 0.639 and b = 0.1296 (Meyers et al.,1992). These measurements were obtained using a continuous flow diffusionchamber (CFDC), whose limited data exhibits roughly a factor of ten higherconcentrations of IN at warmer temperatures than that found with older devicessuch as the filter-processing systems. Recognizing the need to allow verticaland horizontal variations in IN concentrations in mesoscale model simulations,Cotton et al. (2003) modified Eq. (4.20) to include the prognostic variable NIN:

Ni = NIN exp[12.96(Si − 1)], (4.21)

where T < −5 ◦C. The variable NIN can be deduced from the continuous flowdiffusion chamber data and used as a forecast variable in regional simulations(i.e. van den Heever et al., 2006; van den Heever and Cotton, 2007).

As noted above, there is considerable evidence that freezing of drops bycontact nucleation is much more efficient than immersion freezing of drops.


Y

Y

Y

C

C

CC

CC

C

CC

D

D D

D

DVV

V

DC

–20 –15 –10 –5

Temperature (°C)

–25 0

Ice

nucl

ei c

onc

(L–1

)

1

10

100

1000

0.1

10000

FIGURE 4.5 Summary of measurements of contact-freezing ice-nuclei concentrations madeby various authors [Cooper (C), Deshler (D) and Vali (V)]. For comparison, the estimates ofYoung (1974a) (Y) are shown. The regression line is an exponential fit to measurements. Fletcher(1962) ice-nucleus curve is also shown for reference (—). (From Meyers et al. (1992))

Unfortunately, there are no field-deployable devices for measuring contactnuclei concentrations. All that is available are a few exploratory laboratoryexperiments for estimating contact nuclei IN concentrations or the effectivenessof various aerosols to serve as contact nuclei. Using data from vertical windtunnel experiments reported by Blanchard (1957), Young (1974a) deduced thatcontact IN concentrations Nic can be described by

Nic = Na0(270.16Tc)1.3 (4.22)

where Tc is the cloud droplet temperature, and Na0 = 2.0 × 102 liter−1 atsea level. Meyers et al. (1992) fitted data measured by three different prototypedevices simulating contact nucleation by Vali (1974, 1976), Cooper (1980), andDeshler (1982), that are plotted in Fig. 4.5.

Note that those measurements suggest that contact nuclei concentrations aremuch less than Young’s estimates. Meyers et al. (1992) approximated the datawith an equation of the form:

Nic = exp[a = b(273.15Tc)], (4.23)

where a = −2.80 and b = 0.262 when Nic has units of liter−1. Remember thatthe data used to derive Eq. (4.23) are based on a limited number of air samples.


We expect that the concentrations of contact nuclei will vary temporarily andspatially as much as condensation-freezing nuclei do.

Equations (4.19), (4.21) and (4.23) only represent initial ice particleformation on IN and do not necessarily represent actual ice particleconcentrations because other processes such as ice multiplication (seeSection 4.4.2), sedimentation, breakup, and advection can greatly influence theconcentrations of ice particles. Note that Eq. (4.21) allows for both horizontaland vertical variations in IN. Because the aerosol contributing to IN are largeand large aerosols generally decrease with height (Georgi and Kleinjung, 1968;DeMott et al., 2003), we expect that IN concentrations generally decrease withheight as well.

One should not therefore be surprised that many observations of ice particleconcentrations do not show a good correlation with temperature (Gultepe et al.,2001; Field et al., 2005). Gultepe et al. (2001), for example, compiled datafrom the glaciated regions of stratus clouds for a number of field campaignsand found that ice crystals smaller than 1000 µm diameter do not show a goodcorrelation with temperature and that the concentrations of these smaller iceparticles varied up to three orders of magnitude, for a given temperature. On theother hand, measurements of ice particle concentrations in wave clouds whereonly initial ice particles are likely, Cooper and Vali (1981) showed ice particleconcentrations increasing with decreasing temperature.

In summary, we see that there remain many unanswered questionsregarding the concentrations of ice nuclei, their composition, their activationrelative to different environmental factors, and their relationship to ice crystalconcentrations. Work has been hampered by severe difficulties in the precisemeasurement of IN. Moreover, current field deployable devices for measuringIN do not take into account activation of IN by contact freezing. Fletcher’s orMeyers empirical curves are still used and often misused in numerical models,ranging from cloud-resolving models to GCMs. These curves, at best, onlyrepresent the concentrations of ice particles initially formed in clouds andprobably rarely represent ice particle concentrations in most clouds.

4.4.2. Ice Multiplication

As we have seen, it has become increasingly evident that concentrations of icecrystals in real clouds are not always represented by the concentrations of INmeasured or expected to be activated in such environments. In particular, it hasbeen found that at temperatures warmer than −10 ◦C, the concentration of icecrystals can exceed the concentration of IN activated at cloud top temperatureby as much as three or four orders of magnitude (Braham, 1964; Koenig, 1963;Mossop et al., 1970; Mossop and Ono, 1969; Mossop et al., 1967, 1968, 1972;Magono and Lee, 1973; Auer and Marwitz, 1969; Hobbs, 1969, 1974). Theeffect is greatest in clouds with broad drop-size distributions (Koenig, 1963;Mossop et al., 1968, 1972; Hobbs, 1974).


Some explanations or hypotheses that have been proposed to account for thehigh ice particle concentrations observed in some clouds are as follows:

• Ice multiplication by fracturing of fragile ice crystals, which may breakupduring collision with each other. (Vardiman, 1978).• Fragmentation of large drops during freezing. (Mason and Maybank, 1960).• Secondary ice particle formation during ice particle riming. ((Hallett and

Mossop, 1974; Mossop and Hallett, 1974).• Enhanced ice nucleation in the presence of spuriously high supersaturations.

(Hobbs and Rangno, 1985).• Secondary ice particle generation during evaporation of ice particles (Oraltay

and Hallett, 1989; Dong et al., 1994).

Of these processes, the one that has been given the most attention andquantified in models is secondary ice particle formation by the rime-splinterprocess. Laboratory studies by Hallett and Mossop (1974) and Mossop andHallett (1974), confirmed by Goldsmith et al. (1976) have indicated that copiousquantities of splinters are produced during ice particle riming under veryselective conditions. These conditions are:

• Temperature in the range of −3 ◦C to −8 ◦C.• A substantial concentration of large cloud droplets (D > 24 µm)• Large droplets coexisting with small cloud droplets (D < 12.3 µm).

An optimum average splinter production rate of 1 secondary ice particle for250 large droplet collisions occurred at a temperature of −5 ◦C.

This process is consistent with observations of the greatest departure fromIN estimates of ice crystals when clouds contain graupel particles and frozenraindrops, and is consistent with field observations (Hobbs and Cooper, 1987).

There is much indirect or inferential evidence that evaporation enhancesice crystal concentrations. This evidence is perhaps more intriguing than it iscompelling. Some field studies have related unusually high ice nuclei numbers,or unusual increases in ice crystal numbers, to circumstances in which cloudswere evaporating. Cooper (1995), for example, found a 100-fold increase in icecrystal concentrations in the evaporation region of orographic layer clouds. Thelargest ice enhancements in the Cooper study were observed in clouds withtemperatures approaching the onset temperature for homogeneous freezing.Smaller enhancements were found in warmer clouds and no enhancements werefound warmer than about −20 ◦C. Further evidence of the possible role ofevaporation nucleation has been presented by Field et al. (2001) and Cotton andField (2002). They show observational evidence and supporting parcel modelingcalculations from wave cloud studies that suggest ice had to form close to thedownstream edge of the wave cloud. Ice production coincident with the start ofthe liquid cloud, or earlier, would have suppressed the observed liquid cloud.Some of the observations of rapid ice crystal concentration enhancement versus


expected IN concentrations in cumulus cloud studies of Hobbs and Rangno(1985, 1990) and Rangno and Hobbs (1994) were also observed to originate inclose proximity to regions of cloud evaporation. Stith et al. (1994) followed thedevelopment of ice in a cumulus turret near its top at−18 ◦C. During the updraftstages, low ice concentrations were observed in the turret (similar to what wouldbe expected from primary ice nucleation), but during the downdraft stages, theice concentrations increased by an order of magnitude. This observation cannotbe explained by rime splintering.

In summary, it is unlikely that all primary and secondary ice-formingprocesses have been quantitatively identified. Other mechanisms maysometimes operate, but their exact nature remains a mystery. Moreover, ourability to measure small ice crystals, in particular, has significant errors and needimprovement. Consequently, there are large uncertainties associated with ourability to simulate the affect of aerosols on the initiation of ice, and subsequentimpacts on precipitation. This remains as one of the critical problems in cloudphysics.

4.4.3. Ice-Particle Growth by Vapor Deposition

Once ice crystals are nucleated by some mechanism of primary or secondarynucleation, and if the environment is supersaturated with respect to ice, thecrystals can then grow by vapor deposition. Because the saturation vaporpressure with respect to ice is less than the saturation vapor pressure with respectto water, a cloud which is saturated with respect to water will be supersaturatedwith respect to ice. Figure 4.6 shows the variation in supersaturation with respectto ice as a function of temperature for a water-saturated cloud. Note that icecrystals in a water-saturated cloud can experience supersaturation in excessof 10% This leads to a process commonly known as the Bergeron Findeisenmechanism.

If a cloud is at water saturation at −10 ◦C, for example, the supersaturationwith respect to ice will be 10%. As the ice crystals grow by vapor deposition,they deplete the vapor content, thereby driving the environment below watersaturation. The cloud droplets will then evaporate, which helps sustain a vaporpressure difference between ice and water. By this process ice crystals are thensaid to grow at the expense of cloud droplets. This process has been describedby Wegener (1911), Bergeron (1935), and Findeisen (1938).

It should be noted that because the Bergeron-Findeisen process causesthe evaporation of cloud droplets, it can also affect the contact nucleationprocess. Young (1974a,b) and Cotton et al. (1986) have pointed out that if theBergeron-Findeisen process results in the partial evaporation of larger clouddroplets, phoretic scavenging of potential contact nuclei will be enhanced.Thus the Bergeron-Findeisen process and the contact nucleation process canform a positive-feedback loop. Vapor deposition growth of ice Crystals lowersthe saturation ratio below water saturation, causing droplet evaporation. This


Ice supersaturation (%)

wat

er s

ubsa

tura

ted

wat

er s

uper

satu

rate

d

wat

er s

atur

ated

–30

–25

–20

–15

–10

–5

Tem

pera

ture

(°C

)

–35

00 5 10 20 3015 25 35

FIGURE 4.6 Supersaturation with respect to ice as a function of temperature for a water-saturated cloud. The shaded area represents a water-supersaturated cloud.

evaporation favors the phoretic-contact nucIeation of the supercooled dropletsthat have not fully evaporated, and that in turn, grow as ice crystals, furtherlowering the saturation ratio below water saturation, and so on. This processis favored whenever the droplet spectrum is broad enough to allow the partialevaporation of the largest droplets.

It should be noted also that the rate of depletion of supercooled liquid waterby ice crystals growing at the expense of cloud droplets is dependent on theconcentration and size of droplets. In a study of the affects of aerosol pollutionon orographic clouds, Saleeby and Cotton (2005) found that in a pollutedcloud where cloud droplets were numerous and small, ice crystals depletedliquid water much more rapidly than when the cloud was clean and dropletconcentrations were less and droplets were larger. This is simply because, fora given liquid-water content, more numerous small droplets expose a larger netsurface area to the depleted vapor content of the air than a cloud containing few,larger droplets.

Prediction of the vapor deposition growth of ice crystals is complicated bythe fact that ice crystals exhibit differing habits or shapes depending upon thetemperature and supersaturation (with respect to ice) of the environment. Forexample, the results of the laboratory experiments shown in Fig. 4.7 illustratethat the ice crystal habit can change from plates to needles or prisms to platesover less than 1 ◦C [for the definition of the various ice crystal habits, seePruppacher and Klett (1978) and Mason (1971)]. This figure also shows that


plates

plates

plates

dendrites

dendrites

10 cm s–1

10 cm s–1

50 cm s–1

50 cm s–1

equi

axed

pris

ms

disk

s

columns

columns

thin

ner

thic

ker

water

sub

satu

rate

d

wat

er s

uper

satu

rate

dcolumns

needles

–30

–25

–20

–10

–5

Tem

pera

ture

(°C

)

–35

05 10 15 20 25 30

Ice supersaturation (%)0 35

FIGURE 4.7 The shape of a crystal is related to environmental conditions in a complicatedmanner: temperature, supersaturation of the atmosphere with water vapor, and speed offalling all have an effect. Most crystals grow under natural conditions not far removed fromthe diagonal line representing water saturation. The dotted lines to the left of “dendrites” and“needles” show how the speed of falling extends the zones in which those elongated forms grow.(After Keller and Hallett (1982); figure from Hallett (1984))

the ambient temperature and supersaturation are not the only properties thatdetermine the crystal habit. The fall velocity of the crystals and the associatedventilation can also extend the regimes of needle and dendritic forms of crystalsinto regions of subsaturation with respect to water. The habit of growth of anice crystal can have a pronounced influence on the rates of vapor-depositiongrowth. This is especially true for dendritic and needle growth habits, whichgreatly accelerate the deposition growth rate over that of an equivalent sphericalparticle.

Thus, the vapor-deposition (evaporation) equation, Eq. (4.24), for sphericalparticles must be modified to include the role of ice crystal habit. Theconventional approach is to assume that the diffusion of heat and water vaporin the vicinity of a complex-shaped ice crystal behaves in a manner analogousto the rate of electrical charge dissipation from an electrically charged capacitorof similar shape (Jeffreys, 1916). Under this assumption, the rate of mass vapor


a

c

PlateColumnNeedlea

c

a

c

a

a

Spherical

ice particle

Large plate

or dendrite

FIGURE 4.8 Illustration of approximations to crystal shapes by spheroids of revolution.

deposition (sublimation) on an ice crystal can be formulated as

dxi

dt

]V D= 4πCGi (T, P)(Si − 1) f (Re)−

MwLs L f G(T, P)

Ki RaT 2

dxi

dt

]RM,

(4.24)

where xi is the crystal mass, C is the “capacitance” an ice crystal, Si is thesaturation ratio with respect to ice, f (Re) represents a ventilation function ofan ice crystal, and G(T, P) is a thermodynamic function similar to that for waterdrops, but modified to include the saturation vapor pressure and the latent heatbetween vapor and ice. Both the ventilation function and the capacitance varydepending on the particular habit of the ice crystal. The second term on the right-hand side of. Eq. (4.24) represents the contribution to the crystal heat balance bythe latent heat released during riming. Cotton (1970) found that for clouds withliquid-water contents greater than 0.3 g m−3, ice crystal riming (dxi/dt)RMcan warm the crystal sufficiently to drive the surface of the crystal below icesaturation, causing the crystal to sublimate mass. At this point, however, crystalriming growth dominates the growth of the crystal.

The crystal capacitance is generally computed from theoretical electrostaticcapacitance models for simplified shapes such as spheres, disks, and prolate oroblate spheres of revolution. Thus, if we consider a to be the length of the basalplane, and c to be the length of the prism plane of an ice crystal as shown inFig. 4.8, the capacitance can be approximated as follows:for needles,

C = c/ln(4c2/a2); (4.25)

for prismatic columns,

C = ce/[ln |(1+ e)/(1− e)|] where e =√

1− a2/c2 (4.26)


for hexagonal plates,

C = ae/2 sin−1 e, where e =√

1− c2/a2 (4.27)

for thin hexagonal plates or dendrites, the disk approximation

C = a/π (4.28)

may be used, while for spheres,

C = a/2. (4.29)

In order to evaluate the capacitance of each crystal, the bulk geometryof the crystals must be diagnosed or predicted. This is generally done byusing laboratory and/or field data to empirically determine aspect ratios c/aof the crystals or crystal mass versus major dimension (Koenig, 1971; Cotton,1972a; Young, 1974a; Scott and Hobbs, 1977; Cotton et al., 1982; Jayaweera,1971). Alternately, one can follow a more fundamental approach as proposedby Chen and Lamb (1994b). They made use of direct measurements of theindividual growth rates of the prism and basal faces of ice crystals to determinethe individual condensation coefficients of the prism and basal faces. Theyincluded the effects of ventilation of the crystals which enhances the vaporfluxes especially on the prominent edges of the ice crystals. This allows thecrystals to grow to several hundred microns whereupon ventilation forces thecrystals to grow primarily along their long axis.

4.4.4. Riming Growth of Ice Particles

Once ice crystals become large enough, they can settle through a populationof supercooled cloud droplets, colliding and coalescing with them. When theyimpinge upon the ice surface, the droplets immediately freeze, since ice is an“ideal” nucleator. A deposit of frozen droplets called “rime” accumulates on thesurface of the ice crystal. The riming growth process is a collision-coalescenceprocess, analogous to the collision-coalescence growth of liquid cloud droplets.The rate of growth of a single ice crystal of mass xi can be thus described as

dxi

dt

]RM=

∫∞

0A′i (Vi − Vc)E(xi/x)x f (x)dx, (4.30)

where A′i represents the geometric cross-sectional area that an ice crystal sweepsout relative to a cloud droplet of mass x, Vi is the terminal velocity of an icecrystal of mass xi , Vc is the terminal velocity of a cloud droplet of mass x ,E(xi/x) is the collection efficiency between the ice crystal and the cloud dropletx , and f (x) is the spectral density of cloud droplets of mass x ± δx .


Because ice crystals typically fall with their major dimension in thehorizontal plane, the geometric cross section can be approximated as

A′i = (a + r)(c + r)

for needles and columns of dimensions a and c falling through cloud dropletsof radius r , and

A′1 = π(a/2+ r)2

for plates, dendrites, and spherical ice panicles. By selecting a suitable spectraldensity function f (xi ) to keep track of the mass and geometry of ice particles,one can compute the evolution of the ice spectrum by solving quasistochasticintegral-differential growth equations similar to Eq. (4.1) used for clouddroplets. In this manner, Young (1974b) and Scott and Hobbs (1977) havesimulated the evolution of the ice-particle spectrum by storing informationabout particle mass and geometry in a series of continuous “bins.”

If the ice particles are considerably larger than the cloud droplets, then Eq.(4.30) may be simplified by ignoring the cloud droplet radius in evaluatingA′i , by assuming Vi � Vc, and that an average E can be used between allcloud droplets and ice particles of mass xi . Under these assumptions Eq. (4.30)becomes

dxi

dt

]RM= A′i Vi E(i/c)ρ0rc, (4.31)

where rc is the cloud droplet mixing ratio.It should be noted that the ice crystal vapor-deposition growth habits affect

A′i , Vi , and E . Details on how to evaluate the terminal velocity and collectionefficiency of ice crystals collecting cloud droplets can be found in texts byPruppacher and Klett (1978) and Mason (1971) or in papers like Saleeby et al.(2007). As an ice crystal grows by riming, the geometry of the ice crystal ismodified, which, in turn, alters A′i , Vi , and E .

In convective clouds having high liquid-water contents (generally greaterthan 1.0 g m−3) the riming growth of ice crystals can proceed to the pointthat the rime deposit nearly obscures the original crystal habit. Often such aparticle tumbles and rimes to become nearly symmetric or conical in shape. Werefer to such heavily rimed particles as graupel. The density of small graupelparticles may be as low as 0.13 g m−3 (Magono, 1953; Nakaya and Terada,1935). However, as the graupel particle grows larger and falls faster, the densityof the rime deposit may become as high as 0.9 g m−3. It is often thought thatgraupel forms from those ice crystals which have grown for a relatively longtime by vapor deposition and riming. These “large” crystals are candidates forforming graupel. Reinking (1975) has observed that the embryos of graupelparticles are predominantly a select few of relatively smaller crystals that collect


droplets at comparably rapid rates. Perhaps these select few crystals randomlycollect one or more larger than average cloud droplets. The large cloud dropletscould upset the hydrodynamic stability of the crystals, causing them to tumbleand form graupel particles. In contrast, crystals which have undergone longergrowth times by vapor deposition would have large aspect ratios and be lesssusceptible to hydrodynamic instability by the chance collision with a dropletslightly larger in size than the average. Furthermore laboratory experimentsperformed by Fukuta et al. (1982, 1984) suggest that in the temperature range−6 ◦C to −10 ◦C, where more isometrically shaped, columnar forms of vapor-grown crystals prevail, ice particles switch over more readily to the graupelmode of riming growth. It is interesting that the more rapid switchover tothe graupel mode of growth compensates to some degree for the otherwisesuppressed precipitation growth by vapor deposition in that temperature range.

There is also some indication (Holroyd, 1964) that aggregates of ice crystalscan serve as embryos for graupel particles. We will discuss aggregation morefully in the next section. Graupel can also originate as frozen large clouddroplets, drizzle drops, or raindrops. The drops may freeze by contact orimmersion freezing or by collecting a small ice crystal. These large frozendrops have a high density (perhaps greater than 0.9 g cm−3 and therefore rapidlyfall through the population of cloud droplets, collecting them to become high-density graupel particles.

If the cloud contains a high liquid-water content and vigorous updrafts as incumulonimbi, such graupel particles can serve as embryos for hailstone growth.The rate of mass growth of a hailstone may be estimated from the accretionequation, Eq. (4.31), where A′i = πR2

h and Rh is the radius of the hailstone.However, as the hailstone grows by collecting cloud droplets, the latent

heat of freezing is liberated. This latent heat can warm the hailstone to sucha degree that not all the accreted water can freeze. If this occurs, some of theunfrozen water can be shed from the hailstone, thereby limiting its growth(Ludlam, 1951). If the hailstone sheds all unfrozen water, it is said to begrowing in the “wet regime,” and its mass accumulation can be estimatedfrom the thermodynamic budget of the hailstone (Pruppacher and Klett, 1978;Mason, 1971). If all the accreted water can freeze, then Eq. (4.30) representsthe hailstone growth rate and the hailstone is said to be growing in the “dryregime.” There remains, however, some uncertainty as to what fraction of theunfrozen water is actually “shed” from a hailstone growing in the wet regime.Macklin (1961) carried out a set of laboratory experiments simulating thegrowth of hail during the wet regime. He found that much of the unfrozen waterwas not shed in the wake of the accreting object, but instead was incorporatedinto the ice structure as a spongy or mushy ice deposit containing a substantialamount of liquid water. This process is apparently enhanced by the collectionof ice crystals which, Macklin postulated, can form an interwoven mesh ofdendritic crystal structures which trap liquid water. Macklin noted that only attemperatures of −1 ◦C to −2 ◦C was any of the unfrozen accreted liquid shed


in the wake of the simulated hailstones. At colder temperatures, all the unfrozenwater was retained in the “spongy” ice deposit.

4.4.5. Aggregation Growth of Ice Particles

Snowflakes, a common form of precipitation in the wintertime in mid-latitudes,are made up of clusters or aggregates of “pristine” ice crystals. These aggregateshave formed by the collision and coalescence among ice crystals. The collectionprocess can be described by a quasistochastic collection model similar to Eqs(4.1)–(4.4). The problem, however, is complicated by the complex geometriesand orientations of the falling pristine crystals, which affect the formulationof the collection kernel. An additional complication arises from the fact thatonce they have made physical contact, ice crystals do not always coalesce. Inthe case of cloud droplets, there is considerable evidence that the coalescenceefficiency is reasonably high, approaching unity. In this case, the majorproblem in evaluating the collection kernel, Eq. (4.4), is in estimating thehydrodynamic collision efficiency between cloud droplets. This is a ratherstraightforward procedure but by no means a simple problem (Pruppacher andKlett, 1978). In the case of ice crystal aggregation, in addition to the challengingproblem of estimating the hydrodynamic efficiency among ice crystals, wemust also estimate their probability of sticking. Laboratory experiments andinferences from field studies suggest that the coalescence efficiency among icecrystals is higher at warm temperatures (Hallgren and Hosler, 1960; Hoslerand Hallgren, 1960) and is a function of the habit of the ice crystals, withdelicately branched dendrites, having the highest efficiencies (Rogers, 1974).However, Latham and Saunders (1971) found no such temperature dependencein their laboratory experiments. Rauber (1985) has also suggested, from fieldobservational evidence, that a mixture of plane dendrites and spatial dendrites,which have different fall-velocity spectra, favor the formation of aggregates.The exact magnitude of coalescence efficiencies among ice crystals and theirvariation with crystal habit and temperature remains unknown at the presenttime.

Another complicating aspect of snowflake aggregation theory is thatice crystals and aggregates of ice crystals exhibit both horizontal velocityfluctuations and fluctuations in terminal velocity. These velocity fluctuationsare caused by the complex geometry of the crystals, which can produceaerodynamic lifting forces similar to that of an aircraft wing. Further causesof velocity fluctuations include changes in aerodynamic drag forces due tovariations in orientation of the ice crystals and to the shedding of turbulenteddies in the wake of the particle. All of these factors influence the rate ofcollision among ice crystals. For example, when we discussed the coalescenceamong cloud droplets, we defined a collection kernel, Eq. (4.4), which was afunction of the area swept out by the “parent” droplet and the differences inmean terminal velocities. However, in the case of ice crystals, the collection


kernel is influenced by horizontal and vertical velocity fluctuations. Sasyo(1971) examined this effect by laboratory and field experiments as well asnumerical calculations. He applied the kinetic theory of gases to this problemand formulated a collection kernel as

V (ri , r j ) = 2σ√π(ri + r j )

2, (4.32)

where ri and r j represent the effective radii of ice crystals or aggregates, and σis the standard deviation of horizontal and vertical velocity fluctuations. Sasyoestimated a standard deviation of 5 cm s−1 from field experiments. Based onthese results he suggested that collisions of this type would be important duringthe first stages of aggregate formation when ice crystals are of similar size andshape and before a broad spectrum of crystals has formed.

Passarelli and Srivastava (1979) hereafter referred to as PS have also notedthat due to the particular structure of snowflakes, snowflakes of a given masscan have a spectrum of sizes, fall speeds, and shapes. In contrast to thestandard collection kernel, Eq. (4.4), where particles of the same mass have zeroprobability of colliding, such a spectrum leads to a finite probability of collisionof particles of the same mass. PS considered two models of aggregation. Inthe first model snowflakes of a given mass are considered to be spherical andhave a unique diameter but a spectrum of fall speeds. This model bears someresemblance to Sasyo’s model except that horizontal motions of snowflakesare not considered. In the second model snowflakes are also assumed to bespherical, but snowflakes of a given mass are assumed to have a spectrumof bulk densities, which results in a spectrum of diameters and fall speeds.Numerical experiments with the first model suggested that its contributionto snowflake aggregation was only 10% of that due to the standard kernel.However, experiments with the second model, which includes a spectrum ofparticle densities, indicated that the magnitude of the modified kernel is alwaysgreater than that of the standard kernel. Depending on the spectrum width,the modified kernel results in substantially more rapid aggregation. This worksuggests that new approaches to modeling aggregation may be needed. Forexample, if aggregate and pristine ice crystals are only described in a modelas a function of their mass, it may be necessary to develop probability densityfunctions (PDF’s) of their collection kernels for each mass bin.

It is clear from this discussion that there remains a great deal to belearned about snowflake aggregation processes. In addition to the uncertaintiesmentioned above, it should also be noted that aggregation commences from thecollision among pristine crystals whose concentration cannot be consistentlypredicted even within several orders of magnitude of observed values. Thisis especially true at warmer temperatures, where aggregation seems to be themost prevalent. As can be seen from the collection equations [Eqs (4.2) and(4.3)], the aggregation rate is proportional to the product of the concentrationof ice crystals. Thus a threefold order of magnitude uncertainty in ice crystal


concentration results in roughly a sixfold order of magnitude uncertainty inour estimates of the initial rate of aggregation. However, there is considerablemotivation to continue to improve models of the aggregation process, sinceaggregates of dendrites are the most common type of snowfall in mid-latitudes.

4.4.6. Melting of Ice Particles

The melting of ice particles (and associated cooling due to the latent heat offusion), contributes substantially to the formation of thunderstorms (Knupp,1985) and mesoscale downdrafts in tropical squall lines (Houze, 1977) and inwintertime orographic cloud systems (Marwitz, 1983).

The melting of an ice particle is basically a thermodynamic process.Consider a graupel particle of mass Xg which has fallen through the 0 ◦Cisotherm into warmer temperatures. Suppose further that the layer that thegraupel has fallen into contains cloud droplets at the ambient temperature T .Assuming a steady state and that the graupel maintains a surface temperatureT f of 0 ◦C, the rate of latent heat release due to melting must be balanced bythe rate of heat transfer through the layer of water on the graupel surface. Thus,

Lli

[dXg

dt

]melt= −2πDg KT f (Re)(T − T f )− 2πLlvDg Dv f (Re)(ρvs f c)

−

[dXg

dt

]RM

cw(T − T f ), (4.33)

where Dg is the graupel diameter, KT and Dv are the diffusivities for heat andwater vapor, respectively, and cw is the thermal conductivity of liquid water. Thefirst term on the right-hand side of Eq. (4.33) represents the diffusion of heat tothe surface of the melting graupel particle at temperature T f . The second termon the right-hand side of Eq. (4.33) represents the diffusion of water vapor andthe corresponding transfer of latent heat from the graupel surface. A ventilationterm f (Re) is included in both diffusion terms. The third term on the right-hand side of Eq. (4.33) represents the transfer of sensible heat to the graupelas it accretes cloud droplets at the rate (dXg/dt)RM . Mason (1956) consideredthe effects of the first two terms in his study of the melting process. Wisneret al. (1972) and Cotton et al. (1982) have included the last term in their cloudmodels. The transfer of sensible heat during collection can greatly accelerate themelting process, especially in clouds with low cloud bases and correspondingdeep layers of cloud warmer than 0 ◦C.

4.5. SUMMARY OF CLOUD MICROPHYSICAL PROCESSES

In preceding sections, we have seen that the evolution of precipitation in cloudscan take on a variety of forms and involve numerous physical processes. Theevolution of ice-phase precipitation processes is greatly dependent upon the


Cold-basedContinental Clouds

Supply of water vapor

Narrow cloud spectra

frozen dropsheterogeneous

freezing(immersion/contact)

slow broadeningby coalescenceaided by turbulence

continuedcoalescence

DRIZZLEWARMRAIN

condenstion-freezing

vapordeposition

depositionnucleation

of the ice phaseheteorogeneousnucleation

freezing

SLEETGRAUPELS

GRAUPELSSNOW PELLETS

HAILHAIL RAIN SNOW

RAINSLEET WARM

RAINSNOW

GRAINS

freezing

partialmelting

partialmeltingrimed

crystals

clumping

riming

Frozen dropsIce Pellets

Pristine ice crystals

condensation-freezing

icedepositionnucleation

SecondaryIce Particles

melting

shedding

shedding

groupels

secondaryice particles

rimingaggregation

Snow Flakes

SnowCrystals

continuedcoalescence

Drizzle

coalescenceheterogeneousfreezing

(immersion / contact)

Broad cloud droplet spectra

Supply of water vapor

Warm-basedMaritime Clouds

secondaryice particles

FIGURE 4.9 Flow diagram describing microphysical processes, including paths forprecipitation formation. (Adapted from Braham (1968))

prior or concurrent evolution of the liquid-phase. These processes, in turn, aredependent upon the characteristics of the air mass (i.e. aerosols), the liquid-water production of the cloud, the vertical motion of air within the cloud, theturbulent structure, and the time scales of the cloud. We characterize the liquidwater production of the cloud by its base temperature. Clouds having warmbases will have higher cloud base mixing ratios and thus will produce morecondensate than a cold-based cloud over a given depth of cloud. Illustratedin Fig. 4.9 are the different precipitation paths that may occur dependingupon whether the cloud is a cold-based continental cloud versus a warm-basedmaritime cloud. We use the term maritime cloud to represent a very clean airmass and continental to represent one with much higher CCN concentrations.A polluted cloud would have still higher CCN concentrations. The figure doesnot note the speed by which these regimes can produce precipitation. We haveseen that a cloud with a vigorous warm rain process or, what we refer to in thefigure as a warm-based, maritime cloud will produce precipitation much fasterthan a cold-based, continental cloud. The rapidity of glaciation of a warm-basedmaritime cloud is much faster than the cold-based maritime cloud since thepresence of drizzle drops and supercooled raindrops once frozen can rapidlytransform a cloud from an all-water cloud to an ice-dominated cloud. Thisshould not be interpreted to mean that the largest precipitation elements suchas hail would occur in a warm-based maritime cloud. In fact, just the oppositecan take place, as a vigorous precipitation process lower in the cloud can deplete


supercooled water amounts higher up and lower the trajectory of precipitationelements leading to smaller sized hailstones.

This diagram is used to illustrate two distinctly different regimes. Infact, there is a continuum of cloud types between these two extreme states.The heaviest precipitating clouds, and generally the most efficient, are thosethat are warm-based and form in clean, maritime air masses. By contrast,clouds developing in a polluted air mass should be less efficient in producingprecipitation than similar clouds with the same cloud base temperatures. Thissimple reasoning is valid only for single clouds and storms. As we will see, insome cases a suppression or retardation of precipitation in a primary convectivecell could lead to a transformation of a cloud into a longer-lived storm throughthe interaction of cold-pools and in some cases lead to greater amounts ofprecipitation.

4.6. MODELING AND PARAMETERIZATION OF ICE-PHASEMICROPHYSICS

Modeling the microphysics of ice-phase or mixed-phase clouds generallymirrors the approaches used to model liquid-only clouds. The ice-phase,however, is more complicated owing to the departures of ice particles frombeing spheres and to the variations in particle densities. The most advancedmodels which keep track of ice particle habits such as Chen and Lamb (1994a)are so computationally demanding that they are generally limited to the simpleLagrangian parcel type of dynamical frameworks. Single-moment bin-resolvingmodels such as Khain and Sednev (1995) and multi-moment bin models (Reisinet al., 1996a,b) have been extended to include the ice-phase but with simpleparameterizations of ice particle shapes like assuming a single crystal habit, andmaking use of mass-diameter empirical relationships such as that used in bulkmodels. These models have been applied to two and three dimensional cloudmodels (Reisin et al., 1996a,b; Harrington et al., 2000; Khain et al., 2004; Lynnet al., 2005) even though they are very computationally demanding. In fact, a fullbin scheme has been implemented into the Earth Simulator General CirculationModel (GCM)!

Most cloud and storm models, mesoscale models and especially GCMsfollow a bulk microphysics approach that is an extension of the Kessler (1969)warm cloud microphysics philosophy. Like the warm cloud schemes, ice cloudsare assumed to be ice saturated and produce “cloud ice,” the size distributionof all ice hydrometeors is assumed to follow an exponential size-distribution,collection is modeled as a continuos accretion process, and all hydrometeorclasses settle as a mass-weighted mean fall velocity (Cotton, 1972a,b; Lin et al.,1983; Rutledge and Hobbs, 1984). Power-law relationships are generally usedto relate ice particle mass to the dimensions of the ice particles (Locatelli andHobbs, 1974; Lin et al., 1983; Rutledge and Hobbs, 1984; Mitchell et al., 1990;Cotton et al., 2003). We must recognize that there are many differences between


the liquid and ice-phase that make the extension of the Kessler approach tothe ice phase less desirable. For one thing, while supersaturations with respectto water are believed to be less than 1%, that is not the case for the icephase where supersaturations with respect to ice can exceed 10%. Thus whileassuming water clouds remain saturated allowing the diagnosis of cloud watermay not lead to large errors, extension to the ice-phase can lead to muchgreater errors in diagnosing the mixing ratios of vapor grown ice crystals andtheir thermodynamic consequences. Thus Cotton et al. (1986, 2003) replacethe concept of “cloud ice” with predicted pristine ice which retains substantialsupersaturations. Likewise, while modeling collection as a continuous accretionprocess may be valid for graupel particles and hailstones collecting clouddroplets, it is not a good approximation when hailstones collect graupel particlesor raindrops, or aggregates collect pristine ice crystals as the differences interminal velocity between the collector and collectee can be quite small, andthe sizes of both species is appreciable. A somewhat ad hoc approximation wasproposed by Wisner et al. (1972) in which the differences in terminal velocitiesin the collection equation is approximated by the differences in mass-weightedfall speeds. As shown by Verlinde et al. (1990) this can lead to large errors incomputed collection rates.

Due to the storage requirements and computing time required to computedetailed microphysics, explicit prediction of the habits of ice crystals and theirsize distribution have been limited to use in simple one-dimensional, steady-state parcel models (Cotton, 1972b; Young, 1974b, 1975; Chen and Lamb,1994b), at the most in one-dimensional, time-dependent models (Scott andHobbs, 1977). It is generally assumed that the ice particles are either sphericalor have a single crystal habit such as a hexagonal plate. Cotton et al. (2003)allow different crystal habits depending on temperature but these habits are notallowed to advect or diffuse in the cloud or undergo transformations to mixedhabits like capped-columns.

Important to any ice parameterization is the number of ice categories.Because each ice category affects particle fall speeds and collection rates,a larger number of ice categories better represents a cloud composed of,say, vapor-grown ice crystals, partially-rimed ice crystals, aggregates, graupelparticles of varying densities, frozen raindrops, and hailstones of varyingdensities (McCumber et al., 1991). Of course each category of ice addedincreases the computational cost. Even the use of bin-resolving microphysicsmodels does not eliminate the need to select the appropriate number of icecategories. Ferrier et al. (1995) for example, found that using separate graupeland hail categories greatly improved their storm characteristics in comparisonto observations. An example of a large number of ice categories is that of Strakaand Mansell (2005) who decomposed ice into ten categories consisting of icecrystals (three habits), aggregates/snow, graupel (three densities), frozen drops,and hail (small and large).


It has become increasingly popular to expand the number of prognosticparameters or moments describing the basis functions that define the size-distribution of hydrometeors (Nickerson et al., 1986; Ferrier et al., 1995;Meyers et al., 1997; Cotton et al., 2003; Reisner et al., 1998; Morrison et al.,2005a,b). Clark (1974) and more recently Milbrandt and Yau (2005a,b) haveimplemented triple-moment prognostic systems. Milbrandt and Yau appliedtheir triple-moment model to simulating hailstorms in which the sixth-momentor radar reflectivity was the prognostic variable. In addition, while the earlier iceparameterizations were “hard-wired” to the exponential basis function (Cottonet al., 1986; Lin et al., 1983; Rutledge and Hobbs, 1984) more flexible andmore general basis functions such as gamma and log-normal basis functions arenow being used (Ziegler, 1985; Ferrier, 1994; Walko et al., 1995; Clark, 1976;Feingold and Levin, 1986; Feingold et al., 1998).

Finally, as mentioned earlier, in the RAMS model Saleeby et al. (2007)extended the bin-emulating approach to the ice-phase in which realistic (atleast the best available in the literature) collection kernels were used for graupelparticle and hailstone collection of cloud droplets. All these improvements in thebulk parameterization schemes makes them more flexible and closer in characterto the full bin-resolving models.

We now focus briefly on hail parameterization. When we considerparameterization of most hydrometeor species the emphasis is on estimating theaverage water contents and perhaps concentrations. But for hail the emphasisshifts to the tail of the distribution, or largest, most damaging particles. Thusconventional one- or two-moment bulk parameterizations are less suitable.Certainly the most appropriate approach is to use a full quasi-stochastic bin-resolving model. At the time of this writing no one has implemented a fullbin-resolving microphysics model for use in two and three dimensional stormmodels that includes hail thermodynamics (see below). There have been acouple of bin-resolving models of hail growth that have been applied to simpleLagrangian parcel models (Young, 1978; Danielsen et al., 1972). In the pastseveral researchers have implemented a hybrid approach to modeling hail(Farley and Orville, 1986; Johnson et al., 1993). In this approach all non-hail species are modeled following the Kessler-type philosophy. Hail, however,is treated using finite bins with a continuous accretion approximation forhail collecting cloud droplets. Another strategy is to use a bulk microphysicsapproach but to divide the hail spectrum into a small-hail and large-hail modes(Straka and Mansell, 2005). Then there is the bulk approach of Milbrandtand Yau (2005a,b) in which a bulk approach is used but three moments ofthe hail basis function are predicted. The relative advantages of the hybridbulk/bin approach, versus the two-mode bulk approach, versus three-momentbulk approaches have not been quantitatively determined.

One final comment on hail modeling is the thermodynamics of hail growth.Hail collects supercooled drops at such a high rate that the latent heating


of freezing can warm the hailstone to near 0 ◦C. At this point the hailstoneis said to undergo “wet growth” such that only a fraction of the collectedsupercooled water freezes. Some of the unfrozen water can shed, although theamount of water shed can be small (List, 1959, 1960; Macklin, 1961) withsome of it being incorporated into what is called a “spongy hail” mesh (List,1965). Nonetheless determination of whether a hailstone is in wet or drygrowth regimes is important for the shedding calculations and furthermore theradar reflectivity of hailstones is dependent on whether the hailstone surfaceis ice or water coated. A common “textbook” approach to estimating hailstonetemperatures (i.e. Pruppacher and Klett), is to calculate the rate of collection ofsupercooled water by continuous accretion and then calculate the heat budgetof the hailstone by the latent heat of freezing of water and removal of latentand sensible heat by diffusion, and conduction. Normally this is done followingSchumann (1938) and Ludlam (1958) in which the heat storage of the hailstoneis neglected, which permits forming a balance equation between heat releasedand heat absorbed by the hailstone. However, this is a very poor assumption formassive hydrometeors such as hailstones, and probably even graupel particles.To overcome this deficiency, Walko et al. (1995) developed a heat budgetequation for each ice category including hail that retained heat storage. Thiswas done by defining and diagnosing a reference category energy from whichthe fraction of ice vs water can be diagnosed. Such a procedure could be adaptedto a multi-category hail model such as Straka and Mansell’s (2005) model oreven bin-resolving models.

It has become increasingly evident that aerosols can have an importantinfluence on cloud microphysics and especially on precipitation (Levin andCotton, 2007). Most of the studies simulating aerosol/cloud-system interactionshave followed a bin-resolving microphysics approach where not only thehydrometeor spectra are described in discrete bins but so is the aerosol spectrum(Feingold et al., 1999; Yin et al., 2000; Reisin et al., 1996a,b; Khain et al., 1999).The more sophisticated of these models treat CCN as a size-resolved prognosticspecies and track soluble material inside drops (Flossman et al., 1985; Chen andLamb, 1994b; Feingold et al., 1996). This permits not only examining aerosoleffects on clouds but also cloud effects on aerosol.

As far as bulk microphysics models are concerned aerosol effects on cloudshave been mainly examined by varying droplet concentrations which thenimpact on autoconversion processes. This approach can be useful when a singleclass of cloud is represented, such as a thunderstorm or an orographic cloud.Once an ensemble of clouds is represented, or within a model domain layerclouds and deep convective clouds simultaneously occur, this approach ignoresthe role of cloud vertical velocities in determining the concentration of clouddroplets nucleated. Thus Saleeby and Cotton (2005) were motivated to developa parameterization of CCN activation that takes into account not only thevariability of CCN concentrations but also vertical velocity and temperature.First, RAMS had to be changed from a pure diagnostic model of cloud liquid


water using a saturation adjustment approach to predicting the mixing ratio ofcloud liquid water (Walko et al., 2000). Then an ensemble of Lagrangian parcelmodel calculations were performed with the Feingold and Heymsfield (1992)parcel model. Assuming that the aerosol could be chemically characterizedas ammonium sulfate particles and that their size-distribution could bedescribed by a log-normal basis function with a constant breadth parameter, anensemble of realizations was performed, varying temperature, vertical velocity,CCN concentration, and median radius of the CCN distribution. The results ofthose calculations were put into a look-up table that is accessed by RAMS incloud and storms simulations. Furthermore a second cloud droplet mode wasdefined wherein GCCN particles were nucleated assuming they are sodiumchloride particles. Like the more advanced bin-resolving microphysics models,the mass of soluble material was kept track of so that evaporating drops canrecharge the atmosphere with CCN, albeit with an altered size. In summary,the RAMS bulk microphysics model has three prognostic aerosol species, theconcentrations of CCN, GCCN, and IN using Eq. (4.21).

4.7. IMPACT OF CLOUD MICROPHYSICAL PROCESSES ONCLOUD DYNAMICS

What level of complication in the formulation of cloud microphysical processesis needed to simulate a particular cloud system? This question arises naturally,and the answer is influenced, in part, by the scientific objectives. If thegoal is to simulate the average precipitation over a mesoscale region, it maynot be necessary to simulate microphysics in great detail. If, however, it isimportant to distinguish among the various forms of precipitation—such asfreezing rain versus graupel, or aggregates versus pristine or lightly rimed icecrystals—then greater sophistication in the formulation of cloud microphysicsis desirable. The answer to the above question is also driven by the impactof cloud microphysical processes on the dynamics of the cloud system. Ifcloud microphysical processes impact strongly on the dynamics of a cloudsystem, then an incorrect simulation of a cloud microphysical process couldlead, for example, to the simulation of short-lived multicellular thunderstormsin cases where severe, steady supercell storms are actually observed, or weakstorm downdrafts where severe downbursts are observed. Frequently, numericalmodels of clouds experience a bifurcation in dynamic behavior, depending onthe cloud microphysical structure. That is, local changes in cloud microphysicalprocesses can eventually lead to a cloud system having completely differentdynamical characteristics. Let us, therefore, summarize some of the interactionsbetween cloud microphysical processes and cloud dynamics. More detailedexaminations of such interactions will be made in Part II of this volume whenwe study the dynamics of various cloud systems.


4.7.1. Water Loading

In Chapter 2 we noted that the vertical equation of motion for a cloudsystem contains an additional term in the buoyancy term which accounts forthe weight of suspended condensate having combined liquid and ice mixingratios rd and ri , respectively. This term arises from the fact that if clouddroplets and precipitation particles are falling on average at nearly their terminalvelocities, then the sum of the drag forces on a parcel of air due to settlinghydrometeors is equal to the weight of the condensate. As a rule of thumb,roughly 3 g kg−1 of condensate is equivalent to 1 K of negative thermalbuoyancy. One consequence of a precipitation process, therefore, is that itunloads the cloud updraft at higher levels of its condensate and redistributesit to low levels. The additional condensate at lower levels may turn a thermallybuoyant updraft into a downdraft.

4.7.2. Redistribution of Condensed Water into SubsaturatedRegions

A greater consequence of a precipitation process is the redistribution ofcondensation and evaporation processes associated with the condensateredistribution. Numerical experiments by Liu and Orville (1969) and Murrayand Koenig (1972) suggest that the thermodynamic consequences of theprecipitation process are far more important than the water loading effects.

Precipitation may settle into unsaturated air beneath the cloud base whereevaporation commences. The evaporatively chilled air, in turn, can stimulateand intensify downdrafts in the subcloud layer, which can lead to the decay ofthe cloud or contribute to the propagation of the storm as air is lifted along thegust front.

As will be discussed more fully in Chapter 8, The strength of downdraftsand cold pools beneath clouds, in turn, is related to the dryness of the downdraftair as a result of entrainment of dry air in the lower levels of the cloud, theprecipitation rate and to the size-distribution of raindrops. The dependenceon drop size is related to the fact that the surface to volume ratio is greaterfor smaller drops than for larger drops. Thus for a given rainwater content, aprecipitating downdraft composed of smaller drops will evaporate more totalwater when exposed to subsaturated air than one composed of a few largerdrops, leading to colder downdraft air (Hookings, 1965; Knupp, 1985).

4.7.3. Cloud Supersaturation and Cloud Droplet Evaporation

As mentioned earlier it is often assumed that a cloud does not becomesupersaturated with respect to water. This is in accordance with theoreticalstudies (Howell, 1949; Squires, 1952; Mordy, 1959; Neiburger and Chien,1960) and observations (Warner, 1969) in non-precipitating clouds, that peaksupersaturations are generally less than 1% and more typical supersaturations


are between 0.1% and 0.2%. Numerical experiments in a precipitating systemsuggest that there is a much greater tendency for the supersaturation to exceednominal values (Clark, 1973; Young, 1974c). This is a result of the factthat the magnitude of supersaturation is a result of two competing processes:(1) supersaturation production by adiabatic cooling and (2) supersaturationreduction by condensation on cloud droplets. In a non-precipitating cloud,numerous small cloud droplets are able to readily deplete supersaturationas condensation occurs on them. Because precipitation particles form at theexpense of cloud droplets, the net surface area over which condensation takesplace is reduced substantially when water is converted from numerous clouddroplets to fewer, large precipitation elements. As a result, supersaturationsmay exceed 5%, causing a delay in latent heat release compared to situationsin which the cloud’s humidity remains close to 100%.

Perhaps more important is that the concept that clouds do not become highlysupersaturated, motivated a saturation adjustment approach in which cloudwater contents are determined by the difference between total water mixingratios and saturation mixing ratios. The procedure eliminates the need to knowanything about the details of cloud droplet concentrations and sizes. However,detailed bin-resolving microphysics models of Xue and Feingold (2006) andJiang and Feingold (2006) show that increasing concentrations of CCN anddroplets, produced smaller droplets and suppressed drizzle and led to enhancedevaporation of droplets by entrainment. Because, for a given LWC, smallerdroplets evaporate more readily than larger droplets, owing to the greaternet surface area exposed to subsaturated air, entrainment induced evaporativecooling was enhanced when CCN droplet concentrations were high. This ledto greater entrainment rates in clouds, lower cloud liquid-water contents andreduced cloud depth. A model that uses the saturation adjustment approachto determine cloud liquid-water contents would not be able to simulate thisprocess.

4.7.4. Latent Heat Released during Freezing and Sublimation

At levels in the atmosphere colder than 0 ◦C, the potential exists for ice crystalsto be nucleated and grow by vapor deposition and collection of cloud drops.Also, supercooled cloud droplets and raindrops may freeze. The freezing ofcloud droplets and raindrops results in the release of the latent heat of fusion.

The amount of heat liberated is proportional to the amount of supercooledliquid water frozen. Furthermore, as ice crystals grow by vapor deposition, theyrelease the latent heat of sublimation. If the cloud is water saturated and containsa substantial amount of supercooled cloud droplets, however, the full latent heatof sublimation is not absorbed by the cloudy air. This is because ice crystalsgrow by vapor deposition at the expense of cloud droplets, which must evaporateas the saturation vapor pressure is lowered locally below water saturation. Asa result, the evaporating droplets absorb the latent heat of condensation. The


net result of ice crystals growing by vapor deposition and of cloud dropletsevaporating is that the cloud experiences heating only in proportion to the latentheat of fusion (i.e. L f = Ls − Lc, where L f is latent heat of fusion, Ls is latentheat of sublimation, and Lc is latent heat of condensation) and the water massdeposited on the ice crystals. In a glaciated cloud where liquid- water dropletsare absent, a cloud will experience the full latent heat of sublimation as vapor isdeposited on ice crystals.

What this means, of course, is that any cloud will experience an additionalsource of buoyancy as freezing and ice vapor deposition takes place. The latentheating can be rather smoothly released if ice crystals grow by vapor depositionin an air mass that is cooling adiabatically by large-scale lifting. In contrast, itcan take place as a burst of energy release in convective towers if large quantitiesof supercooled water suddenly freeze.

It is important to recognize that ice-phase-related latent heating beginsto become important at levels in the atmosphere where the latent heatof condensation is greatly reduced. This is due to the fact that at coldertemperatures the vertical variation of saturation vapor pressure with respectto water diminishes substantially as a parcel of air is cooled by adiabaticexpansion. Ice-phase latent heating is also important in disturbed environmentssuch as tropical cyclones (Lord et al., 1984) and mesoscale convective systems(Chen and Cotton, 1988), where the environmental sounding is nearly wetadiabatic. In such an environment the cloud realizes little buoyancy gain fromthe latent heat of condensation, while ice-phase latent heating can contribute tosubstantial convective instability. We will discuss the importance of ice-phaselatent heating more fully in Chapters 7–9.

4.7.5. Cooling by Melting

Melting is distinctly different from the freezing process, in which freezingis distributed through a considerable vertical depth. By contrast, cooling bymelting is quite localized. In the case of the more stratiform precipitation,melting can result in a well-defined isothermal layer. The stability of thisisothermal layer can inhibit the downward penetration of upper troposphericwinds to lower tropospheric levels.

As will be discussed more fully in Chapter 8, cooling by melting can bean appreciable contributor to thunderstorm downdraft strengths (Knupp, 1985).Furthermore, as shown by van den Heever and Cotton (2004) the size of meltinghydrometeors is important to the cooling rate of storm downdrafts. This isbecause melting is a result of heat exchanges at the surface of the ice particle.Thus, like evaporation of raindrops, a precipitating downdraft composed of alarge number of small hailstones will melt faster than one with the same amountof condensed mass but composed of a few hailstones. We will discuss theconsequences of this on the dynamics of thunderstorms in Chapter 8.


4.7.6. Radiative Heating/Cooling

We shall see in the next chapter that cloud radiative heating and cooling rates arestrongly modulated by the microphysical structure of a cloud. The reflection ofincident solar radiation at the tops of clouds is dependent upon the concentrationand size of cloud droplets or the concentration, size, and habit of ice crystals.The absorption and heating by solar radiation, in turn, is a function primarilyof the integrated liquid-water path and secondarily of the size of the droplets.Likewise, the absorption of terrestrial radiation is related to the integrated liquid-water path. As a consequence, both the rates of solar and terrestrial radiativecooling are primarily related to the vertical distribution of condensate in a cloud.

Because precipitation processes alter the vertical distribution of conden-sate, they also strongly impact upon longwave and shortwave radiative heat-ing/cooling rates in clouds. The transformation of a cloud from the liquid phaseto the ice phase can impact the radiative properties of a cloud. Furthermore, theparticular habit of ice crystal growth can affect the rate of absorption of solarradiation which, in turn, can alter the thermodynamic stability of the cloud sys-tem. We will examine these processes more quantitatively in the next chapter aswell as in Chapter 6 and Chapter 10.

4.7.7. Electrical Effects

The influence of cloud electrification processes on the dynamics of clouds isdiscussed more fully in Chapter 8. Here we will simply enumerate some of thehypothesized ways in which cloud electrification can affect cloud dynamics:

(1) Localized heating arising from lightning discharges.(2) Levitation of cloud particles, which alters the terminal velocity of particles,

causing a redistribution of condensate.(3) Enhancement of droplet and ice-particle coalescence, thus enhancing

precipitation formation in a redistribution of condensate.

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