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SANTACHIARA ET AL.
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The Mystery of Ice Crystal Multiplication in a Laboratory Experiment
GIANNI SANTACHIARA, FRANCO BELOSI, AND FRANCO PRODI
Institute of Atmospheric Sciences and Climate (ISAC), National Research Council, Bologna, Italy
(Manuscript received 18 April 2013, in final form 29 August 2013)
ABSTRACT
This paper addresses the problem of the large discrepancies between ice crystal concentrations in clouds
and the number of ice nuclei in nearby clear air reported in published papers. Such discrepancies cannot
always be explained, even by taking into account both primary and secondary ice formation processes. A
laboratory experiment was performed in a cylindrical column placed in a cold room at atmospheric pressure
and temperature in the 2128 to 2148C range. Supercooled droplets were nucleated in the column, in the
absence of aerosol ice nuclei, by injecting ice crystals generated outside in a small syringe. A rapid increase in
the ice crystal concentration was observed in the absence of any known ice multiplication. The ratio between
the mean number of ice crystals in the column, after complete droplet vaporization, and the number of ice
crystals introduced in the column was about 10:1. The presence of small ice crystals (introduced at the top of
the column) in the unstable system (supercooled droplets) appears to trigger the transformation in the whole
supercooled liquid cloud. A possible explanation could be that the rapidly evaporating droplets cool sufficiently to determine a homogeneous nucleation.
1. Introduction
Clouds in Earth’s atmosphere modify climate, as they
interact with both incoming shortwave and outgoing
longwave radiation (Stephens 2005; Gettelman et al.
2012). More than 50% of Earth’s precipitation originates in the ice phase (Heymsfield et al. 2006). Clouds
are usually classified as ‘‘warm’’ (liquid water droplets),
‘‘cold’’ (ice clouds), or ‘‘mixed phase’’ (the latter consisting of supercooled liquid droplets and ice particles).
Ice nucleation in clouds is one of the key processes in
initiating precipitation (Heymsfield et al. 2006).
The formation of ice in clouds can occur in two ways:
(i) primary processes (nucleation of ice from the liquid
or water vapor phases), either homogeneously or heterogeneously triggered by aerosol particles called ice
nuclei (IN), and (ii) secondary processes. The homogeneous nucleation process involves only pure water or
solution droplets. Micrometer-sized pure water droplets
should nucleate homogeneously in stationary conditions
at about 2388C (Pruppacher and Klett 1997). Homogeneous nucleation from nearly pure water does occur in
Corresponding author address: Gianni Santachiara, Institute of
Atmospheric Sciences and Climate (ISAC), National Research
Council, Via Gobetti, 101, Bologna 40129, Italy.
E-mail: [email protected]
DOI: 10.1175/JAS-D-13-0117.1
Ó 2014 American Meteorological Society
the atmosphere—for example, in the updrafts of cumulus clouds (Rosenfeld and Woodley 2000; Heymsfield
et al. 2005). The probability of a given droplet freezing
spontaneously increases with its volume, so the largest
droplets tend to freeze first and the smallest tend to
freeze last.
Four heterogeneous nucleation mechanisms are distinguished for atmospheric ice formation: deposition,
condensation freezing, contact freezing, and immersion
freezing. In addition to these standard model nucleation
mechanisms, Durant and Shaw (2005) generalized the
notion of contact nucleation to include the crystallization from particles contacting a supercooled droplet
from the inside out, as well as from the outside in. It has
been proposed that this may occur while a drop evaporates. However, such an association has not been found
in the Ice in Clouds Experiment–Layer Clouds (ICE-L)
field program (Heymsfield et al. 2011). Heterogeneous
freezing occurs at lower supersaturation and higher
temperatures T than homogeneous freezing.
Secondary ice formation processes can occur through
1) fracture of ice crystals exposed to dry air layers or
collision of preexisting ice crystal (Vardiman 1978;
Oraltay and Hallett 1989);
2) fragmentation of large individual cloud drops during
freezing in free fall, thus playing a role in multiplying
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the number of ice particles in clouds (Hobbs and
Alkezweeny 1968); and
3) fragmentation of freezing droplets following their
collision with ice particles in the cloud (riming),
which is known as the Hallett–Mossop process
(Hallett and Mossop 1974).
For the past four decades, heterogeneous ice nucleation has been parameterized as two independent
functions of temperature (Cooper 1986; Fletcher 1962)
and of supersaturation over ice or water (e.g., Huffman
1973; Meyers et al. 1992). There have also been several
experimental attempts to represent concentrations of
IN as a function of temperature and supersaturation
(e.g., Berezinsky and Stepanov 1986; Cotton et al. 1986).
Recently, several different heterogeneous ice nucleation parameterizations have been suggested that take
into account that different types of aerosols can act as IN
with different nucleating efficiency (Diehl and Wurzler
2004; Khvorostyanov and Curry 2004; Phillips et al. 2008;
Connolly et al. 2009; DeMott et al. 2010).
A few published papers report ice crystal number
concentrations in good agreement with the number
concentrations of ice nuclei in orographic wave clouds,
where secondary ice formation processes are avoided
(Cooper and Vali 1981; Eidhammer et al. 2010). Some
authors have even found a rapid formation of exceptionally high ice particle concentrations at cloud temperatures between 248 and 2108C (Koenig 1963; Mossop
et al. 1968; Hallett et al. 1978; Hobbs and Rangno 1990;
Rangno and Hobbs 1991; Choularton et al. 2008). Many
published papers show large discrepancies between ice
crystal concentrations in clouds and the number of IN in
nearby clear air (Koenig 1963; Mossop et al. 1968; Hobbs
1969; Mason 1973; Hobbs 1974; Hallett et al. 1978; Langer
et al. 1979; Keller and Sax 1981). Such large discrepancies
can sometimes be explained by considering either primary (homogeneous and heterogeneous) or secondary
ice formation processes (Bower et al. 1996), but in other
cases the cause of the disagreement is unclear (e.g.,
Prenni et al. 2007).
Rosinski and Morgan (1991) suggested that, in convective clouds, evaporating droplets leave behind ‘‘dry’’
residues, which may act as IN activated either by condensation followed by freezing or by deposition, depending on relative humidity. Field et al. (2001) and
Cotton and Field (2002) also found a greatly enhanced
ice nucleation in the wave-cloud evaporation zone.
The observation of ice nucleation in the downdraft
phase of the wave cloud led Cooper (1995) to suggest that evaporating droplets would cool rapidly—
perhaps enough to enhance the homogeneous freezing
rate.
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In most previous measurements the strong increase in
IN concentrations with increasing supersaturation with
respect to ice has not been considered sufficiently, along
with the time lag for ice germs to grow to ice crystals
(Isaac and Douglas 1972).
In natural clouds, however, there are circumstances
that can lead to supersaturation with respect to water Sw
considerably in excess of zero. It has been suggested
(Gagin 1972; Nix and Fukuta 1974; Gagin and Nozyce
1984; Fukuta and Lee 1986; Rangno and Hobbs 1991)
that the increased ice-nucleating ability of particles at
high Sw might explain ice enhancement in some clouds.
Additional explanations for the observed discrepancy
between IN and ice particles in clouds may certainly
hinge on some deficiencies in counting the number of
IN. In fact, IN may represent less than 1 in 106 of the
aerosol population, presenting a difficult measurement
challenge. It is now recognized that many circumstances
result in erroneous measurements of small ice crystals
because ice crystals larger than a few hundred microns
can impact the probe’s upstream tips or inlet of the
measuring instruments and shatter into small fragments
(Field et al. 2006; Jensen et al. 2009; Korolev et al. 2011).
Artifact crystal production can be extreme in some circumstances, reaching an enhancement factor of 100
(DeMott et al. 2011).
The occurrence of a thin layer of supercooled liquid
water drops at the top of both stratiform and convective ice-phase clouds has been frequently observed
(Cunningham 1951; Cooper and Vali 1981; Hobbs and
Rangno 1985; Rauber and Grant 1986; Rauber and
Tokay 1991).
Westbrook and Illingworth (2011), using 4 years of
radar and lidar observations of layer clouds, showed that
supercooled liquid water occurs at the top of the majority of ice-cloud layers warmer than 2278C, and is
almost always present when the cloud-top temperature
is greater than 2208C. The ice nucleation proceeds via
the liquid phase at temperatures greater than 2208C,
and deposition nucleation plays a minor role in the
formation of ice at these temperatures. Ansmann et al.
(2009) studied the formation of the ice phase in tropical
altocumulus and found that the liquid droplets were
always observed first, before the appearance of ice falling beneath, and concluded that ice nucleation mainly
starts via the freezing of supercooled droplets. The
process of droplet and ice formation was observed at
cloud-top temperature lower than 2108C. Similarly, de
Boer et al. (2011), from ground-based lidar, radar, and
microwave radiometer observations of low- and midlevel
Arctic stratiform clouds (,4-km altitude), observed that
clouds composed entirely of ice occur less frequently than
liquid-topped mixed-phase clouds at temperatures warmer
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SANTACHIARA ET AL.
than 2258 to 2308C. These results indicate that ice formation generally occurs in conjunction with liquid at these
temperatures and suggest the importance of the liquiddependent ice-nucleation mechanism at T . 2258C.
Hobbs and Rangno (1985, 1990) and Rangno and
Hobbs (1991) documented numerous instances of ice
concentrations exceeding those that could be explained
by primary nucleation (homogeneous and heterogeneous),
even when the Hallett–Mossop riming–splintering process could not have produced ice concentrations of the
magnitude that they observed in the time available.
They gave a concise outline of their rationale in Hobbs
and Rangno (1998). Their most salient conclusions were
that the clouds glaciated much faster than the Hallett–
Mossop theory could explain, the crystal habits were
often not compatible with the temperature range in
which the Hallett–Mossop mechanism operates, and
high concentrations of small ice particles appeared
concurrently with frozen drizzle drops, and not afterward, as would be expected if the smaller crystals were
a product of riming–splintering. The authors concluded
the final paper in the series with a summary of their
findings (Rangno and Hobbs 1994). They admitted that
the mechanism giving rise to the high concentrations of
ice that they report was perplexing, closing with the
statement ‘‘In summary, the origin of ice in cumuliform
clouds remains a mystery.’’ This can be considered a
conclusion similar to that of Vali (2004), who recognized
that ‘‘there is ample evidence for unexpectedly high ice
concentrations in many situations where the Hallett–
Mossop mechanism is not active.’’
Heymsfield and Miloshevich (1993) and Baker and
Lawson (2006) found in wave clouds, from either allwater or mixed phase, a sudden reduction in the number
concentration of small water particles, accompanied by
an increase in the mean particle size and ice concentration. They also concluded that ‘‘most aspects of the
observations are consistent with basic cloud physics, but
some aspects remain difficult to interpret.’’ Similarly,
Cantrell and Heymsfield (2005) stated ‘‘The Hallett–
Mossop process does not explain some documented
cases of secondary production. Is there another process
operating, or is the Hallett–Mossop process not sufficiently understood?’’ In addition, Fridlind et al. (2007)
wrote ‘‘The lack of correlation between ice particle
number concentration and temperature points to the
likelihood of a ‘universal mechanism’ for which no explanation currently exists, although evaporation nuclei
and evaporation freezing are candidates.’’
The aim of the performed experiments is to provide a
contribution toward explaining the discrepancies observed between in-cloud ice crystal concentrations and
the number of IN.
91
FIG. 1. A schematic diagram of the experimental apparatus (not
scaled). Holes open only during droplet introduction (1), ice crystal
introduction (2), and ice crystal and droplet measurements (3).
2. Experimental
The experimental instrumentation consisted of a
Plexiglas cylinder (diameter of 12 cm and height of
140 cm) placed in a cold room (Fig. 1). It was operated
at atmospheric pressure and at temperatures in the
2128 to 2148C range, as this is the interval showing the
highest difference between the vapor pressure of the
water and of the ice. Liquid droplets of Milli-Q water
(Millipore Corporation) were produced using an ultrasonic nebulizer (Aerosol Project–Artsana). The droplets were injected at the top of the column for a time of
90 s. During this process an opening at the bottom of
the column allowed the gas and the droplets to exit.
Therefore, the column was completely filled and droplets became supercooled in a few seconds.
A similar ultrasonic nebulizer injected droplets into
a syringe (volume of 120 cm3). The droplets were subsequently frozen by introducing for a few seconds a
small wire cooled with liquid nitrogen, generating small
ice crystals in the syringe. In a second step, the crystals
were injected into the column full of supercooled
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JOURNAL OF THE ATMOSPHERIC SCIENCES
FIG. 2. Average size distribution of the droplets in the column.
droplets. In a time of the order of 25 s the droplets disappeared, and only crystals remained in the column.
Crystals grew at the expense of water vapor from evaporating supercooled droplets and fell to the bottom of
the column. The total number of ice crystal concentrations in the column and syringe were obtained by measuring the ice crystal concentrations with an optical
particle counter (OPC, Grimm Model 1.108, Grimm
Aerosol Technik, GmbH). The OPC measures the light
elastically scattered from a single particle illuminated by
a well-defined light source while it is passing through the
sensing volume of the instrument. The scattered light
intensity is utilized as a measure of the particle’s size.
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The sampling flow rate is 1.2 L min21 with a size separation range of 0.3–20 mm (15 channels). The instrument
is widely used in environmental aerosol monitoring
(Grimm and Eatough 2009). The ice crystal counting
was considered from 0.5 mm since the first size bin
channel might overcount the particles, as shown in a recent comparison work (Belosi et al. 2013). The upper
limit of size bin counting is given by the sampler inlet
efficiency. Following the Agarwal and Liu criteria for
sampling from still air (Agarwal and Liu 1980), the efficiency can be considered unitary until about 35 mm.
After nucleation, ice crystals generated in the column
were sampled with the replica technique described by
Schaefer (1956). Pieces of microscope slides (;1 cm2)
fixed on stubs and covered with a thin layer of 2% formvar
solution in chloroform were exposed to the falling
crystals at the bottom of the column. The slides were
then placed inside a desiccator, where crystals and chloroform evaporated. Subsequently, the stubs were examined by SEM.
In additional runs, maintaining the same experimental
conditions (time of droplet injection, etc.) the size distribution of the droplets in the column was measured
with the same OPC (Fig. 2). Sampling of droplets was
also performed using the same procedure followed for
the crystals (Fig. 3).
The relative humidity before droplet nucleation was
measured by a dewpoint hygrometer (Hygro-M2,
FIG. 3. Replica of droplets on slides covered with formvar.
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SANTACHIARA ET AL.
FIG. 4. Average size distribution of the ice crystals in the column
(ICc) and in the syringe (ICs).
General Eastern Instrument, United States). The gas
phase was found to be saturated with respect to water
before the nucleation process. The liquid water content
(LWC) of the cloud in the column was evaluated in two
ways—that is, from the size distribution measured with
the OPC, and by aspirating all the supercooled droplets
from the column onto an inertial impactor located at the
bottom of the column. This device has a small container
at the bottom that is filled with liquid nitrogen. Therefore,
the supercooled droplets that impact the surface of the
impactor freeze instantaneously. The formed ice deposit
was then transferred into a small plastic cylinder and
weighed. The LWC turned out to be in the range of 1.2–
2.5 g m23. Both methods agree in the arithmetic droplet
diameter and the LWC values.
To test the possible influence of the walls on the nucleation process, experiments were performed by filling
the column with droplets, but without introducing crystals. As we did not observe nucleation of the droplets, the
nucleation due to column walls can be excluded.
3. Results and discussion
Figure 4 shows the average size distribution of the ice
crystals generated in a syringe outside the column (ICs)
and the ice crystals in the column (ICc) after the complete droplet vaporization. The total number of ice
crystals in the column and in the syringe are reported
starting from 0.5 mm in size. Although the size distributions have to be considered qualitative, owing to the
nonspherical shape of the ice particles, the counts should
be considered accurate.
The number of the crystals generated in the syringe
and introduced at the top of the column turns out to be
lower than the number of crystals measured at the base
of the cloud chamber in all of the size ranges considered,
the difference increasing with increasing crystal diameter. This means that the crystals inside the column
grow at the expense of water vapor supplied by the
93
FIG. 5. Mean, maximum, and minimum values of ICc and ICs.
evaporating droplets, linked to the fact that vapor
pressure over ice eice is less than vapor pressure over
liquid water (Wegener–Bergeron–Findeisen process).
When ice crystals were not introduced in the column by
the syringe, the supercooled droplets remained liquid
and settled.
Figure 5 shows the averaged total number of ice crystals in the column, after the complete droplet evaporation, and the number of small ice crystals introduced at
the top of the column. The figure shows also the minimum and maximum values obtained in all the experiments. Considering all the experiments, the mean number
of ice crystals in the column, after complete droplet
vaporization, is found to be higher than the mean
number of ice crystals introduced into the column. The
ratio is about 10:1. The increase in the number of ice
crystals measured in the column, compared to those introduced through the syringe, is not due to the presence
of ice nuclei particles. In fact, preliminary measurements
inside the column with a condensation particle counter
(CPC, TSI Model 3775) gave a very low aerosol particle
number concentration. Therefore, aerosol ice nuclei were
absent in the column and heterogeneous nucleation
processes due to aerosol particles should be excluded. In
addition, the measured time necessary to obtain a column completely filled by crystals (about 25 s), after the
introduction of ice crystals from the syringe, is lower
than the sedimentation time of the crystals from the top
of the column, where nucleating ice crystals were introduced. It should be noted that even the largest crystals sampled at the bottom of the column (about 50 mm
in diameter) should have a terminal velocity of about
4 cm s21 and a falling time of about 35 s (Heymsfield
1972). Ice crystal fallout was complete within 2–4 min.
The nucleation process appears to be much faster than
the growth and subsequent fallout of the crystals.
In our experimental runs, the secondary mechanism
of ice generation should be excluded, since the riming
process is absent, owing to the lack of turbulence and the
fact that droplets are prevalently less than 20 mm (mean
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FIG. 6. Replica of the ice crystals captured on glass slides.
diameters of 9.9 and 7.8 mm, respectively, counting the
droplets on microscope slides or considering the size
distribution measured with OPC). The image of crystals
sampled at the bottom of the column, after complete
droplet evaporation, confirms this statement. Ice crystals were prevalently hexagonal plates and did not show
any breaking on tips (Fig. 6). We performed in addition
experiments at T 5 258C (i.e., at a temperature where
the Hallett–Mossop effect could be effective), but no
increase in the crystal concentration was observed in the
column compared to values in the syringe.
The experiments undertaken seem to show that the
thermodynamically unstable system considered (supercooled droplets) promptly passed to a more stable state
(Murray et al. 2008). The presence of small ice crystals
(introduced at the top of the column) in the unstable
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SANTACHIARA ET AL.
system appears to trigger the nucleation process in the
whole supercooled liquid cloud. Despite the wide variety of techniques used in the study of homogeneous
freezing, relatively few consider the freezing of evaporating water droplets (Kuhns and Mason 1968; Kr€
amer
et al. 1996; Shaw and Lamb 1999).
A possible explanation of our laboratory results could
be that evaporating droplets rapidly cool sufficiently to
freeze, even at higher air temperatures (about 2138C)
than those usually reported for homogeneous nucleation
in wave clouds (e.g., Heymsfield and Miloshevich 1993;
Field et al. 2001). The great decrease in droplet temperature during evaporation could depend on the fact
that, in the interaction mechanism between droplets and
ice crystals, a stationary field is not expected to form.
Fuchs (1959) showed that the thermal relaxation time of
the droplet in evaporation (tth 5 r2/a, where r is the
droplet radius and a is the thermal diffusivity) is much
longer than the relaxation time of the water vapor diffusion field (t vd 5 r 2/D, where D is the water vapor
diffusion coefficient). Therefore, a marked decrease in
droplet temperature can occur during a fast evaporation.
This conclusion agrees with the results of Satoh et al.
(2002), who studied the cooling–freezing phenomena of
a water droplet of pure distilled water due to evaporation in an evacuated chamber. The results showed that
the water droplet is cooled by the evaporation of water
itself and that freezing starts from its surface. The
droplet froze within a few seconds through a remarkable
supercooling state, and the cooling rate of the water
droplets was dominated by heat transfer within the
droplet. Similar conclusions were shown by Murty and
Murty (1972) and Jung et al. (2012), who obtained
a homogeneous nucleation of evaporating droplets at
the free surface (instead of heterogeneous nucleation
due to the presence of the substrate) at the initial temperature of 2158C.
The hypothesis that, even in a simple system like pure
water or nitric acid solution, the ice nucleation may
occur at the air–droplet interface, and that the homogeneous crystallization may indeed be a surface- and not
a volume-related rate process, was theoretically and
experimentally suggested by Djikaev et al. (2002) and
Tabazadeh et al. (2002). This conclusion goes against the
classical theory of homogeneous crystallization, where
freezing is assumed to initiate inside a droplet and to
propagate afterward to the surface.
A second mode of ice formation suggested by
D. L. Mitchell (Desert Research Institute, 2013, personal
communication) is as follows. The very high liquid water
content implies a high cloud droplet concentration. As
terminal velocity of droplets and crystals are similar, ice
crystals introduced in the column and droplets can be
in close proximity for a relatively long period of time,
where a strong vapor flux from the droplet to the crystals
is manifested. During such an event, the most energetic
water molecules in the droplet are likely to evaporate
first, leaving lower-energy molecules behind. The lowerenergy molecules are more likely to be hydrogen bonded
in a hexagonal geometry similar to that of hexagonal
ice. This hexagonal orientation in the liquid phase is likely
to trigger a freezing event at these initial temperatures
(2128 to 2148C), thus freezing the droplet, which then
grows into an ice crystal.
In conclusion, it is our opinion that the performed
laboratory experiment provides useful information on
the problem of the large discrepancies between ice
crystal concentrations in clouds and the number of ice
nuclei in nearby clear air reported in published papers.
Additionally, it offers support to field campaigns, where
a thin layer of supercooled liquid water drops at the top
of both stratiform and convective ice-phase clouds is
reported, and ice nucleation mainly starts via the freezing
of supercooled droplets, even at cloud-top temperatures
of up to 2108C (Cooper and Vali 1981; Ansmann et al.
2009; de Boer et al. 2011; Westbrook and Illingworth
2011).
Acknowledgments. We are grateful to David L. Mitchell
(Desert Research Institute, Reno, Nevada) for helpful
suggestions and two anonymous reviewers.
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