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Polar Science 5 (2011) 399e410www.elsevier.com/locate/polar

Comparative analysis of measurements of stratospheric aerosol bylidar and aerosol sonde above Ny-Alesund in the winter of 1995

[Comparative analysis of lidar and OPC observations]

Koichi Shiraishi a,*, Masahiko Hayashi a, Motowo Fujiwara a, Takashi Shibata b,Masaharu Watanabe c, Yasunobu Iwasaka d, Roland Neuber e, Takashi Yamanouchi f

aFaculty of Science, Fukuoka University, Nanakuma, Jonan-ku, Fukuoka 814-0180, JapanbGraduate School of Environmental Studies, Nagoya University, Furo-cho, Chikusa-ku, Nagoya 464-8601, Japan

cAdvanced Industrial Science and Technology (AIST), Kansai, Midorigaoka Ikeda, Osaka 563-8577, Japand Institute of Nature and Environmental Technology, Kanazawa University, Kakuma Kanazawa, Ishikawa 920-1192, Japan

eAlfred Wegener Institute for Polar and Marine Research, D-14473 Potsdam, GermanyfNational Institute of Polar Research, Midoricho, Tachikawa, Tokyo 190-8518, Japan

Received 21 December 2009; revised 4 July 2011; accepted 22 August 2011

Abstract

Solid polar stratospheric cloud (PSC) layers observed by lidar and a balloon-borne optical particle counter (OPC) on 17December 1995 are reexamined in a comparative analysis framework. The typical radius of solid particles in the observed PSC isdetermined through the comparative analysis to have been approximately 2.3 mm. A backward trajectory analysis for the air mass inwhich the solid particles were observed shows that the air mass had experienced temperatures 2e3 K below the frost point of nitricacid tri-hydrate (NAT) during the 4 days preceding the observations. The back-trajectory analysis traces the air mass back tonorthern Greenland and Ellesmere Island on 16 December, one day before the observations. A microphysical box model is used toinvestigate possible mechanisms of formation for the observed solid particles. The results of this model suggest that the solidparticles formed under mesoscale temperature fluctuations associated with mountain lee wave activity induced by the relativelyhigh terrestrial elevations of northern Greenland and Ellesmere Island.� 2011 Elsevier B.V. and NIPR. All rights reserved.

Keywords: Polar stratospheric clouds; Stratospheric aerosol; Lidar; Aerosol sonde

1. Introduction

Polar stratospheric clouds (PSCs) play a well-knownrole in stratospheric ozone depletion in polar regions.PSCs serve as sites for heterogeneous chemical

* Corresponding author.

E-mail address: siraisi@fukuoka-u.ac.jp (K. Shiraishi).

1873-9652/$ - see front matter � 2011 Elsevier B.V. and NIPR. All rights

doi:10.1016/j.polar.2011.08.003

activation of halogen compounds, converting thesespecies into potentially ozone-destroying radicals.

PSCs are classified into two types: Type I and Type II.A Type II PSC appears at temperatures below the frostpoint of ice (Tice) and is composed of ice particles (Pooleand McCormick, 1988). A Type I PSC appears attemperatures above Tice but below 196 K. The lattertemperature corresponds roughly to the frost point of

reserved.

400 K. Shiraishi et al. / Polar Science 5 (2011) 399e410

nitric acid tri-hydrate (TNAT). Type I PSCs are furtherclassified into Type Ia and Type Ib. A Type Ia PSCconsists of solid particles composed of nitric acidhydrates such as nitric acid tri-hydrate (NAT) and nitricacid di-hydrate (NAD) (Iraci et al., 1994; Marti andMauersberger, 1993; Worsnop et al., 1993). A Type IbPSC consists of liquid particles composed of a super-cooled ternary solution (STS), which contains H2SO4,HNO3, and H2O (e.g., Carslaw et al., 1997).

The Arctic stratosphere, though cold, is warmer thanthe Antarctic stratosphere because the Arctic polarvortex is more unstable than the Antarctic polar vortex.This additional instability in the Arctic polar vortexresults from atmospheric wave activity, which is moredisruptive in the northern hemisphere than in thesouthern hemisphere. In Arctic winter, most observedPSCs are of Type I (Massoli et al., 2006; Maturilli et al.,2005). As temperature in the Arctic stratospheredecreases close to the frost point of ice, the backgroundsulfate aerosols grow larger, absorbing ambient atmo-spheric vapors (such aswater vapor and nitric acid vapor)and forming STS. Comparisons of field observationswith model simulations (e.g., Saitoh et al., 2002;Tabazadeh et al., 1994) and direct in situ observations(Schreiner et al., 1999) both show that the liquid particlesformed by this process are composed of STS. Solidparticles, on the other hand,werefirst detected in direct insitu observations of theArctic stratosphere byVoigt et al.(2000), who found NAT particles with diameters lessthan 2 mm in nacreous clouds. NAT is thought to be moststable in the lower stratosphere (Hanson andMauersberger, 1988), but it is widely believed thatdirect nucleation of NAT is difficult because ofa substantial nucleation barrier (Prenni et al., 1998;Worsnop et al., 1993). Most hypotheses suggest that iceor meta-stable hydrates, such as NAD, are producedinitially by cooling associated with mesoscale fluctua-tions, such as mountain lee waves, and are only latertransformed to NAT (Carslaw et al., 1999; Iraci et al.,1998; Meilinger et al., 1995; Tabazadeh and Toon,1996; Tsias et al., 1997); however, no nitric acidhydrates other than NAT have ever been detected by insitu field observations.

Field observations and statistical analyses of gravitywaves suggest that solid PSCs may form undertemperature fluctuations induced by inertia-gravitywaves (Shibata et al., 2003; Yoshiki and Sato, 2000).Sulfate aerosols containing meteoritic smoke particlesmay also contribute to the formation of solid particles(Bogdan et al., 2003; Curtius et al., 2005). Despite thesepresentations of viable formation processes, however,the basic mechanisms affecting the growth process and

composition of solid PSCs still remain uncertain.Measurements of particle sizes and concentrationswithin solid PSCs in a variety of meteorological situa-tions are required to improve understanding of thetypical formation process and composition of thesePSCs.

We have observed stratospheric aerosols using bothlidar and balloon-borne optical particle counter (OPC)instruments above Ny-Alesund (79�N, 12�E), Norwayevery winter since 1994. These observations monitorPSC activity under cold stratospheric conditions. In thisstudy,we reexamine observationsmade on 17December1995, when we detected the presence of a solid PSC.Analyses of the lidar and OPC observation have beenreported previously (Hayashi et al., 1998; Shibata et al.,1999; Shiraishi et al., 2003; Watanabe et al., 2004), butthese previous analyses did not consider the formationprocess of this PSC. In this paper, we present a compar-ative analysis of lidar and OPC observations of solidPSCs and discuss possible formation processes of thesolid PSC detected on 17 December 1995. For the latter,we refer to recently presented hypotheses regardingformation processes for solid PSCs.

2. Measurement techniques

The lidar measurements were taken using a NeodiumYAG laser with two wavelengths: 1064 nm and 532 nm.The backscattered light was collected by a 35-cmdiameter telescope and split into three detector chan-nels. Two of these channels detected backscatteringparallel and perpendicular to the polarization plane of thetransmitted laser beam at 532 nm. The third channeldetected backscattering signal at the fundamentalwavelength (1064 nm). Signal pulses for each channelwere measured using a photon counting technique. Thelidar system is described in more detail by Shibata et al.(1997).

The lidar data is analyzed using three parameters: thescattering ratio R, the aerosol depolarization ratio dM,and the Angstrom coefficient a. These three parametersare derived using an iterative method (Shibata et al.,1996). R is defined as (bRþbM)/bR, where bR and bMare the molecular and aerosol backscattering coeffi-cients, respectively. R-1 is often roughly proportional tothe mass mixing ratio of aerosols (Pinnick et al., 1976).dM is defined as bM,t/(bM,tþbM,jj), where bM,t andbM,jj are the aerosol backscattering coefficientsperpendicular and parallel to the polarization plane ofthe transmitted laser beam at 532 nm, respectively, suchthat bM,tþbM,jj is equal to bM. The parameter dMindicates the nonsphericity of particles. The height

401K. Shiraishi et al. / Polar Science 5 (2011) 399e410

profiles of dM are normalized under the assumption thatthe depolarization ratio of atmospheric molecules is0.05 (Adachi et al., 2001). For particles that arecompletely spherical, such as liquid droplets, dM is equalto zero. High values of dM indicate the presence of non-spherical particles, such as crystallized particles. a isdefined as eln(bM1064/bM532)/ln(1064/532), wherebM1064 and bM532 are the aerosol backscattering coeffi-cients for the laser wavelengths of 1064 nm and 532 nm,respectively. a represents the contribution of the aerosolparticle size distribution under the assumption that bMdepends on the wavelength l according to bMfl�a.

Data from 72,000 laser pulses was integrated toproduce a backscatter profilewith a vertical resolution of900 m at a time resolution of approximately 2 h.Measurement uncertainties in scattering ratio and aero-sol depolarization ratio were obtained from the erroranalysis conducted by Russell et al. (1979). In theabsence of PSCs, the uncertainties in scattering ratiobelow 25 km are less than 2% for the 532 nmwavelengthand less than 15% for the 1064 nm wavelength. Theuncertainty in the aerosol depolarization ratio below25 km is less than 15%.

The particle size distribution data used in this anal-ysis was collected by a balloon-borne optical particlecounter (OPC). The OPC was originally designed forballoon-borne observations of the stratosphere by theatmospheric aerosol research group at Nagoya Univer-sity and the engineering group at SIGMATEC (Hayashiet al., 1998; Tsuchiya et al., 1996; Watanabe et al.,1997). The OPC counts and measures the size ofaerosol particles in ambient air that is drawn into itsscattering chamber. The size of particles is determinedfrom the intensity of scattered light, using Mie theoryand assuming particles to be spherical with refractiveindex of 1.4. These assumed parameters correspond to70 wt% sulfuric acid droplets. The concentrations ofparticles with radii between 0.15 mm and 1.8 mm aremeasured separately for each of five size channels(r > 0.15, 0.25, 0.4, 0.6, and 1.8 mm). The typicaluncertainty in OPC measurements is 20%e30%(Watanabe et al., 2004). The statistical uncertainty inOPC measurements is estimated as the standard error ofthe mean, assuming a Poisson distribution. All OPCdata used in this paper are averaged vertically over 1 kmin altitude. The balloon sonde typically ascends at a rateof approximately 4e5 m s�1. The sampling timeinterval for the OPC is 20 s, corresponding to a verticalresolution of 80e100 m. Approximately 10 rawobservations are included in each vertically averageddata point. The air sampling rate is approximately5.0 � 10�5 m3 s�1. The minimum particle

concentration detectable by the OPC is 100 m�3 in all ofthe five size classes.

3. Observations: solid PSC observed on 17December 1995

In the winter of 1995e1996, the Arctic polar vortexwas unusually stable and extremely low temperatureswere often detected in the Arctic stratosphere (Naujokatand Pawson, 1996). The temperature in the lowerstratosphere above Ny-Alesund was intermittentlycolder than the PSC threshold from early Decemberuntil the end of February, and the lidar observationsfrequently detected PSCs during this period (Shiraishiet al., 1997). Both lidar and OPC observations wereperformed on 17 December 1995. The balloon carryingthe OPC was launched at 15 UTC at a location 2 km tothe north of the lidar station. Solid PSCs were detectedby both the lidar and the OPC.

Fig. 1 shows vertical profiles of the scattering ratio(Fig. 1a), the aerosol depolarization ratio (Fig. 1b), andthe Angstrom coefficient (Fig. 1c) obtained from lidarobservations. Profiles of temperature, TNAT, and Ticeare shown in Fig. 1d (Hanson and Mauersberger, 1988;Marti and Mauersberger, 1993). The profiles of frostpoint temperatures were calculated assuming differentconcentrations of H2O (between 4 and 5 ppmv) andHNO3 (between 1 and 15 ppbv). The temperatureprofile was observed by the OPC.

On 17 December 1995, a broad background aerosollayer existed at altitudes between 9 and 30 km, witha peak scattering ratio of 1.1 at 14 km. Sulfuric aerosollayers that originated in the June 1991 eruption of Mt.Pinatubo were still detected during the winter of 1995.A PSC layer composed mainly of solid particles wassuperposed on this broad background layer, at altitudesbetween 19 and 24 km. Both the scattering ratio andthe aerosol depolarization ratio were enhanced in thisPSC layer, as shown in Fig. 1. The Angstrom coeffi-cient was smaller in the PSC layer relative to thebackground aerosol layer between 12 and 15 km,suggesting that the PSC layer contained large solidparticles, such as NAT and/or NAD. The temperatureof the PSC layer was colder than TNAT, assumingmixing ratios of 5 ppmv H2O and 10 ppbv HNO3.

Fig. 2 shows the time progression of vertical profilesof scattering ratio R and aerosol depolarization ratio dMobtained on 17 December 1995. Values of R and dMchanged gradually in the PSC layer. Daily rawinsondeobservations conducted by the Alfred Wegener Instituteindicate that wind speeds at altitudes of 20e23 km overNy-Alesund were approximately 20 m s�1 at the time of

Fig. 1. Vertical profiles of (a) scattering ratio, (b) depolarization ratio, (c) Angstrom coefficient, and (d) temperature observed on 17 December

1995. Profiles of the frost points of nitric acid tri-hydrate (TNAT) and ice (Tice) are also shown in panel (d) as green thin and blue thick dashed lines,

respectively. TNAT and Tice are calculated under varying assumptions of tracer amounts within the ranges 4e5 ppmv of H2O (WXX: XX � 0.1

ppmv of H2O) and 1e15 ppbv of HNO3 (NYY: YY ppbv of HNO3).

402 K. Shiraishi et al. / Polar Science 5 (2011) 399e410

balloon launch. The observed time variations of R anddM in the PSC layer over Ny-Alesund are consistent withthe spatial variation of a PSC with a horizontal extent of1000 km. During the OPC sounding (from 15 UTC to 17UTC), no significant time variation was observed in thevertical profiles of R and dM.

Ten day isentropic backward trajectories are calcu-lated using the NIPRTrajectoryModel and Japanese 25-year Reanalysis Project data (Tomikawa and Sato, 2005;http://www.firp-nitram.nipr.ac.jp/en/), and are used toinvestigate the temperature history of the solid PSCdetected on 17 December 1995. The trajectories arecalculated for the air mass at 21.5 km at 16 UTC. Theuncertainty of the trajectory-derived air mass history isassessed by initializing four additional trajectories atlocations �1� in latitude and longitude relative to Ny-Alesund.

Fig. 3 shows the temporal histories of temperature(Fig. 3a), the saturation ratios of liquid particles overNAT (SNAT) and NAD (SNAD) particles (Fig. 3b), and the

Fig. 2. Time variation of observed vertical profiles of (a) scattering

ratio and (b) depolarization ratio on 17 December 1995.

terrestrial height of the trajectory-estimated location ofthe air mass (Fig. 3c) during the 10 days prior toobservation of the PSC overNy-Alesund. The frost pointtemperatures of NAT (TNAT) and H2O (Tice) are alsoshown in Fig. 3a, along with the temperature of STSformation (TSTS). TSTS was approximated as TNAT-3.6 K(Biele et al., 2001). SNATand SNADwere calculated usingthe method introduced by Salcedo et al. (2001). The air

Fig. 3. Temporal histories of (a) temperature (black solid lines) and

the frost point temperatures of nitric acid tri-hydrate (TNAT) (black

broken line), super-cooled ternary substances (TSTS) (gray broken

line), and water (Tice) (gray solid line); (b) saturation ratios of liquid

particles over nitric acid tri-hydrate (SNAT) (black lines) and nitric

acid di-hydrate (SNAD) (gray lines) particles, and (c) terrestrial height

of the location of the air mass that arrived at 21.5 km altitude over

Ny-Alesund at 15 UT on 17 December 1995.

403K. Shiraishi et al. / Polar Science 5 (2011) 399e410

mass experienced the temperatures colder than TNAT

during the 4 days preceding the observation, but neverexperienced temperatures colder than TSTS (Fig. 3a).The trajectory analysis indicates that the air mass passedover northern Greenland and Ellesmere Island(approximately 1500 m elevation) during the 1e2 dayspreceding the observation (Fig. 3c).

In addition to the history of the air mass, we estimatethe trajectory of the balloon-borne OPC during ascentusing the Japanese 25-year Reanalysis data. When theballoon reached 21 km altitude (at approximately 16UTC), its estimated location is approximately 50 kmsouth-east of the lidar station. The temperature history ofthe air mass detected by the OPC at 21 km is not shown,but it is similar to the temperature histories estimated forthe air mass detected by the lidar and shown in Fig. 3a.The optical properties detected by the lidar in the PSClayer did not vary significantly during the OPC obser-vation (Fig. 2). We assume, therefore, that the opticalproperties of the observed PSC did not vary significantlyduring the observation period.

4. Discussion

Solid PSCs were detected by both OPC and lidar on17 December 1995. To discuss the formation process ofthe observed PSC, wemust estimate the size distributionof PSC particles. To accomplish this, we fit a bimodallognormal distribution function to the OPC data. TheOPC used in the present observation measures radii ofaerosol particles in 5 size classes between 0.15 and1.8 mm. Five values are insufficient for fitting toa bimodal size distribution function, so we assume thatthe total number concentration of the first mode N1 is20 cm�3. This assumption is based on condensationnuclei (CN)measurementsmade inside theArctic vortexin the absence of PSCs byDeshler andOltmans (1998) atKiruna, Sweden on 28 January 1995, who observeda particle concentration of 20 cm�3 at approximately20 km.Using this size distribution potentially introducesadditional biases, such as instrument uncertainty andaerosol variability; however, uncertainties in the totalnumber concentration of the firstmode have little impacton our estimations of backscattering coefficients andradii of the second mode at typical PSC altitudes.

We estimate the size distribution function by mini-mizing the following parameter t:

t¼ 1� dblidar

532

blidar532

!� �blidar

532 � bOPC532

�2þ 1� dblidar

1064

blidar1064

!

� �blidar1064 � bOPC

1064

�2 ð1Þ

where blidar532 and blidar1064 are the backscattering coefficientsobserved by the lidar at wavelengths of 532 and1064 nm, respectively, dblidar532 and dblidar1064 are the uncer-tainties of backscattering coefficient at both wave-lengths, and bOPC532 and bOPC1064 are the backscatteringcoefficients obtained by fitting the OPC data.

For our estimate of the backscattering coefficients forbackground aerosols (temperatures higher than 196 K),the aerosol composition and refractive index werederived using techniques introduced by Steele andHamill (1981). The particle size distribution observedby the OPC is bimodal at the typical altitude of PSCs(21e24 km), where an enhancement of the depolariza-tion ratio was detected by the lidar. For PSCs, aerosolparticles of the first mode of the size distributions areassumed to be ternary solution droplets composed ofH2O, HNO3 and H2SO4, and aerosol particles of thesecond mode are assumed to be composed of NAT. Thecomposition of ternary solution droplets is estimatedusing the box model introduced by Carslaw et al. (1995)and the refractive index is derived using the methodpresented by Krieger et al. (2000). The refractive indexof NAT is assumed to be 1.50.

The aerosol backscattering coefficients were calcu-lated from the derived size distribution functions byapplying Mie theory under the assumption that allaerosols detected by the OPC were spherical particles.These calculations assume mixing ratios of 10 ppbvHNO3 and 5 ppmv H2O, which are consistent withsatellite and field observations (Deshler and Oltmans,1998; Santee et al., 1996; Vomel et al., 1997).

Fig. 4 shows the aerosol size distribution functionsobtained for 1-km intervals between 15 and24kmaltitudeon 17 December 1995. The fitting procedure using thebimodal size distribution function yields reasonableresults. Between 15 and 20 km, the derived size distribu-tions are nearly monomodal, and do not exhibit a distinctsecond (large particle) mode. Between 21 and 25 km,where the lidar detected an enhancement of the depolar-ization ratio, the size distributions are distinctly bimodalwith greater concentrations of large particles in the secondmode. The median radius of particles in the second modeis 2.3 mm between 21 and 22 km. All particles includingsolid particles are assumed spherical in our calculations.For a more detailed analysis, it will be necessary toaccount for scattering by non-spherical particles.

Assuming that PSCs detected by the lidar consist ofan external mixture of liquid (STS) and solid (NAT)particles, and assuming that the depolarization ratios ofeach type of particles are constant, the total backscat-tering coefficients of STS (bSTS) and NAT (bNAT)particles can be expressed as

Fig. 4. Aerosol size distributions at altitudes between 15 and 24 km obtained by fitting with optical particle counter (OPC) measurements on 17

December 1995. Solid line denotes the integrated particle number concentration. Dotted line shows the differential particle size distribution. Solid

black circles denote the integrated particle concentrations obtained by OPC and the error bars are the uncertainties in OPC measurements.

404 K. Shiraishi et al. / Polar Science 5 (2011) 399e410

bSTS ¼fdNAT � dair �RðdNAT � dÞg

dSTS � dNATbR and ; bNAT

¼ fdSTS � dair �RðdSTS � dÞgdNAT � dSTS

bR

where dSTS and dNAT are the depolarization ratio ofSTS and NAT particles, respectively (Daneva et al.,2003). The contribution of bNAT to the total backscat-tering coefficient can then be derived as bNAT/(bNAT þ bSTS). We estimate the backscattering coef-ficients of STS and NAT particles under the assump-tions that dSTS ¼ 0 and dNAT ¼ 0.35 (Adachi et al.,2001). Fig. 5 shows vertical profiles of the aerosolbackscattering coefficients observed by the lidar and

estimated from OPC and lidar data at wavelengths of532 nm (Fig. 5a) and 1064 nm (Fig. 5b). Fig. 5c showsvertical profiles of the ratio bNAT/(bNAT þ bSTS) esti-mated by lidar data and the ratio of backscatteringfrom the second mode particles to the total backscat-tering coefficient (first mode þ second mode) in theheight range of 19e24 km. These two ratios agree wellat the altitude of the observed PSC (21e24 km). Thisagreement suggests that solid NAT particles representa substantial component of the second particle sizemode. The concentration and median radius of solidPSC particles detected by the lidar and OPC between21 km and 22 km are estimated to be 6.0 � 10�4 cm�3

and 2.3 mm, respectively.

Fig. 5. Aerosol backscattering coefficients at wavelengths of (a) 532 nm and (b) 1064 nm derived from the lidar (solid line) and estimated from

the size distributions fitted for the aerosol layers observed by the lidar and optical particle counter (OPC) (filled circles) on 17 December 1995.

The error bars denote the uncertainties in backscattering coefficients derived from the lidar. (c) Vertical profiles of the contribution of solid polar

stratospheric clouds to the total 532 nm backscattering coefficient observed by the lidar, bNAT/(bNAT þ bSTS) (solid line) and the contribution of

the second mode of the size distribution to the total 532 nm backscattering coefficient observed by the OPC, bmode2/(bmode1 þ bmode2) (filled

circles).

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The trajectory-derived temperature history of thePSC air mass indicates that this air mass experiencedtemperatures lower than the frost point of NAT duringthe 4 days preceding the observation (Fig. 3a). The airmass passed over northern Greenland and Ellesmereapproximately 1e2 days prior to the observation. Bothof these locations are characterized by relatively hightopography (w1500 m). The solid particles observedby the lidar and OPC may therefore be influenced bymesoscale temperature fluctuations induced by moun-tain lee waves.

We consider two mechanisms by which the solidparticles detected by the lidar and OPC may be formed:homogeneous nucleation under synoptic temperatureconditions and particle formation influenced by meso-scale temperature fluctuations induced by mountain leewaves.We examine these two candidate processes usingthe microphysical box model introduced by Daneva andShibata (2001). This model estimates the temporalevolution of PSCs along an isentropic temperaturehistory (as in Fig. 3a). The growth of individual particlesis estimated by calculating the condensation of ambientvapors, such as nitric acid and water vapor, onto theparticle surface. Daneva and Shibata (2001) describe themodel in detail.

The microphysical box model includes a volume-based parameterization of homogeneous nucleation ofNAD particles. The production rate P of NAD is

estimated as J�V, where J and Vare the nucleation ratecoefficient and the total volume of liquid particlessaturated with respect to NAD, respectively. Recentlaboratory measurements suggest that stratosphericconditions may favor the formation of meta-stable NADparticles (rather than NAT) because the potential barrierto NAD nucleation may be lower (e.g., Natsheh et al.,2006). For simplicity, this model assumes that thenucleation rate coefficient is constant. NADparticles areinitially formed from liquid particles under saturatedcondition (SNAD> 1). Themodel converts newly formedNAD particles into NAT particles instantaneously. Eachsimulated particle ensemble grows or shrinks accordingto the exchange of mass between the gas phase and thecondensed phase.

The model simulation is initialized with zero PSCparticles 120 h prior to the observations (Fig. 3). Thesize distribution of the background aerosols is initial-ized as the first mode of the bimodal size distributionobserved by the OPC at 21e22 km. The numberconcentration of solid particles is assumed to benegligible at 21e22 km because of the low value of thesecond mode relative to the first mode. Sulfuric acidaerosols are divided into 30 size bins according toradius, from a minimum radius of 0.01 mm toa maximum radius of 0.1 mm. The each size bincontains particles of uniform radius and composition.The initial gas mixing ratios of HNO3 and H2O are

406 K. Shiraishi et al. / Polar Science 5 (2011) 399e410

assumed to be 10 ppbv and 5 ppmv, respectively. Thetemporal evolutions of the STS and NAD particleensembles are simulated along the temperature historyuntil the time of observation.

Fig. 6 shows backscattering coefficients estimatedfrom the STS and NAT particle ensembles for a varietyof NAD nucleation rate coefficients. Backscatteringcoefficients for the NAT particle ensemble are esti-mated assuming that all NAT particles are sphericalwith a refractive index of 1.5. The size distribution ofliquid aerosols does not deviate substantially from theinitial size distribution during this model simulationbecause the air mass does not experience temperatureslower than TSTS. The backscattering coefficient forliquid aerosols also stays constant over the prescribedtemperature history. The value of the backscatteringcoefficient estimated for the NAT particle ensemble(2.5 � 10�9 sr�1 m�1) agrees well with that of solidparticles estimated from the lidar observations whenthe NAD nucleation rate coefficient is prescribed as2.5 � 104 cm3 s�1.

Prescribing the NAD nucleation rate coefficient as2.5� 104 cm3 s�1 results in hydrate particle productionat a rate of 1.4 � 10�6 cm�3(air) h�1. This value isapproximately 4 orders of magnitude higher than thatsuggested by recent laboratory experiments (Knopfet al., 2002; Mohler et al., 2006). Larsen et al. (2004)observed that solid PSCs detected by their balloon-borne instruments seemed to have formed by homoge-neous nucleation under synoptic temperature condition.Using a PSC box model with surface-based homoge-neous nucleation of nitric hydrate particles to matchtheir field observations, Larsen et al. (2004) estimated

Fig. 6. Model-simulated backscattering coefficients for the super-

cooled ternary solution (STS) (filled circles) and nitric acid tri-

hydrate (NAT) (open squares) particle ensembles calculated using

homogeneous nucleation rate coefficients estimated at the times of

lidar and optical particle counter (OPC) observations.

a nitric hydrate particle production rate of2.5e3.6 � 10�5 cm�3(air) h�1, one order of magnitudehigher than ours. The temperature history of the solidPSCs detected by Larsen et al. (2004) showed that thesePSCs had experienced high saturation ratios of NAT attemperatures approaching the frost point of ice. Thesolid PSC detected over Ny-Alesund on 17 December1995 did not experience such cold temperatures duringthe ten days prior to observation (Fig. 3a), nor did liquidaerosols in the air mass experience such high saturationratios of NAT and NAD (10 > SNAT and 3 > SNAD)during that period. Mohler et al. (2006) measuredvolume-based homogeneous nucleation rates of NAD ina super-cooled binary HNO3/H2O solution, and reportedNAD formation rates of 10�6 cm�3 h�1 at temperaturesbetween 185 K and 190 K and SNAD of approximately8e9. They further suggested that homogeneous NADnucleation could only play a role in the formation ofPSCs at values of SNAD that approach or exceed suchvalues.

Fig. 7 shows the model-simulated size distributionsof the STS and NAT particle ensembles at completion ofthemodel run, which corresponds to the time of our lidarand OPC observations. The STS particle ensemble sizedistribution does not change substantially from theinitial prescribed size distribution. The NAT particleensemble size distribution broadens considerably, on theother hand, with a pronounced increase in the numberconcentration of larger (micron-sized) particles. Larsenet al. (2004) reported similar changes in the size distri-bution of their simulated solid particle ensemble. Thisbroad distribution does not match the OPC observations

Fig. 7. Model-simulated size distributions for the super-cooled

ternary solution (STS) (open circles) and nitric acid tri-hydrate

(NAT) (filled circles) particle ensembles at the times of lidar and

optical particle counter (OPC) observations. The simulation incor-

porates the homogeneous freezing process of nitric acid di-hydrate

(NAD) over the temperature history shown in Fig. 3.

Fig. 8. Model-simulated size distributions for the super-cooled

ternary colution (STS) (open circles) and nitric acid tri-hydrate

(NAT) (filled circles) particle ensembles at the times of lidar and

optical particle counter (OPC) observations. NAT particles are

assumed to have formed during the 28 h preceding the observations,

as the air mass passed over the high altitude (1500 m) locations of

north Greenland and Ellesmere Island.

407K. Shiraishi et al. / Polar Science 5 (2011) 399e410

at 21e22 km, however. The OPC data indicate a narrowdistribution that differs substantially from the model-simulated distributions. Considering the differencesbetween the simulated and observed size distributionsand the relatively low values of SNAD experienced by theair mass during the days immediately preceding theobservational time, it appears unlikely that the solid PSC

Fig. 9. Response of t, referring to Eq. (1), to variations in the median radi

between 21 and 22 km.

particles detected on 17December 1995were formed byhomogeneous nucleation of NAD at synoptictemperatures.

We also consider the possibility that the formation ofsolid PSC particles was influenced by mountain leewaves. For simplicity, we assume that the solid NATparticles formed immediately during mesoscaletemperature fluctuations that occurred as the air masspassed over high altitude (w1500 m) location approxi-mately 28 h prior to the observation (Fig. 3c).

As above, we estimate the temporal evolution ofPSC along the isentropic temperature history using themicrophysical box model formulated by Daneva andShibata (2001); however, in this calculation, 0.03%of STS particles (6.0 � 10�4 cm�3) in all particle sizebins are assumed to freeze and form NAT as the airmass passes over high altitude terrain, 28 h prior to theobservation. The value 0.03 corresponds to the ratio ofthe concentration of particles in the second mode to thetotal concentration of particles (first mode þ secondmode). No new NAT particles are assumed to formbefore completion of the model run at the time ofobservation.

Fig. 8 shows the model-simulated size distributionsof the STS and NAT particle ensembles at the time ofobservation. The width of the simulated NAT sizedistribution is very narrow: almost all simulated NATparticles have a radius of approximately 3.3 mm. The

us and width of the second mode of the size distribution at altitudes

408 K. Shiraishi et al. / Polar Science 5 (2011) 399e410

simulated NAT size distribution broadens only slightlyif NAT particles are assumed to form during a timewindow spanning 24e40 h prior to the observationrather than only 28 h prior, with radii ranging from 3.1to 3.6 mm.

The size distributions obtained using the lidar andOPC observations correspond more closely to themodel simulations that assume particle formationoccurs in mountain lee waves rather than the modelsimulations that assume particle formation occurs insynoptic temperature conditions. Our estimated sizedistributions, in which the size distribution function ischosen to minimize the parameter t, suggest that themedian radius of solid particles (second mode) at21e22 km altitude is approximately 2.3 mm. Thisvalue is outside of the range of particle sizes that theOPC instrument can detect. Fig. 9 shows the responseof t to variations in the width and median radius of thesecond mode of the size distribution function. Forvalues of the median radius between 2.1 and 3.1 mm, tis minimized for narrow distribution widths, between1.01 and 1.30. The lidar and OPC observations pre-sented here therefore suggest that solid PSCs arestrongly influenced by mountain lee waves. Our lidarstation (Ny-Alesund) is located west of Spitsbergenand does not neighbor any mountains higher than 1 kmaltitude. Some studies using field observations suggestthat air masses passing over Ny-Alesund are not typi-cally influenced by mountain lee waves (e.g., Bieleet al., 2001); however, Greenland is only approxi-mately 700 km west of our station. Carslaw et al.(1999) investigated the occurrence of solid particleformation in mountain lee waves using trajectoryanalysis and a mountain wave model, and identified thearea over eastern Greenland as an important locationfor solid particle formation induced by mountain leewave activity. Air masses that pass over Greenlandoften proceed downstream to Ny-Alesund; therefore,many PSCs detected over Ny-Aalesund might beinfluenced by mountain lee waves.

5. Conclusions

A comparative analysis of lidar and OPC observa-tions of a solid PSC over Ny-Alesund on 17 December1995 is performed to investigate the formation processesof solid PSCs. Themedian radius of solid particles in theobserved PSC is estimated to be approximately 2.3 mm.Backward trajectory analysis for the air mass containingthe solid particles shows that the air mass had experi-enced temperatures 2e3 K below the frost point of NATduring the 4 days prior to the observations. The

trajectory analysis traces the air mass back to northernGreenland and Ellesmere Island on 16 December, oneday before the observations. We examined two candi-date processes for the formation and growth of solid PSCparticles in the stratosphere: homogeneous nucleationunder synoptic temperature conditions and homoge-neous nucleation in mesoscale temperature fluctuationsinduced by mountain lee waves. The results of a micro-physical box model simulation indicate that nucleationunder the influence of mountain lee waves is consistentwith the observations, while nucleation under synoptictemperature conditions is not. Analysis of the sizedistribution of solid particles derived from both lidar andOPC observations strengthens this conclusion, sug-gesting that solid particles in this PSC were likelyformed under mountain lee wave activity occurring overnorthern Greenland and Ellesmere Island.

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