# Global Distributions and Changes in Stratospheric Particles

Sophie Godin, Lamont Poole, Slimane Bekki, Terry Deshler, Niels Larsen,

Thomas Peter

► To cite this version:

Sophie Godin, Lamont Poole, Slimane Bekki, Terry Deshler, Niels Larsen, et al.. Global Distributions and Changes in Stratospheric Particles. Scientific Assessment of Ozone Depletion: 1998, 44, World Meteorological Organization, pp.3.1-3.40, 1999, GORMP. hal-03325693

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CHAPTER 3

# Global Distributions and Changes in Stratospheric Particles

**Lead Authors:**

S. Godin
L.R. Poole

**Coauthors:**

S. Bekki
T. Deshler
N. Larsen
T. Peter

**Contributors:**

A. Adriani
J. Barnes
R. Bevilacqua
C. David
G. Di Donfrancesco
M. Fromm
M. Hervig
H. Jäger
B. Luo
R. Neuber
M. Osborn
W. Renger
L. Thomason
O. Uchino
M. Wirth
SCIENTIFIC SUMMARY

Much progress has been made recently in our understanding of the two major classes of stratospheric particles: stratospheric sulfate aerosol (SSA), and polar stratospheric clouds (PSCs). Thermodynamic models have provided a clearer picture of particle behavior at low temperatures, while a richer and longer measurement suite has increased our knowledge of particle formation processes, the dispersal and decay of volcanic SSA, and particle climatology.

* There is no clear trend in background SSA from 1979 to 1997. SSA levels in late 1997 were below those observed before the 1991 Mt. Pinatubo eruption and are likely still decreasing. Hence, any anthropogenic contribution to the SSA layer must be smaller than previously estimated from observed changes from 1979 to 1989. Peak aerosol scattering ratios in 1997 were about 40% greater than those observed during 1979, but due to uncertainties and natural variability in the measurements, this difference must be viewed with caution at present.

* It is not clear that the 1979 minimum SSA period was truly free of volcanic influence. Recent model calculations of SSA production from known tropospheric sulfur sources significantly underestimate the 1979 observations. Other non-volcanic sources are thought to be insignificant.

* Post-volcanic SSA decay varies with time, space, and aerosol property. The $e^{-1}$ decay time for column backscatter following the eruption of Mt. Pinatubo was about 1 year until 1994, and nearly twice as long (1.8 years) from 1994 to 1997. Derived surface areas decayed back to pre-Pinatubo levels in about 3.5 years at 25 km and about 5 years at 15 km. Surface area decayed 20-30% more slowly than backscatter or mass.

* PSC observations are still divided into two broad classes: Type 1 PSCs, containing nitric acid ($HNO_3$) as a major component, that form at temperatures above the water ($H_2O$) ice point; and Type 2 PSCs, containing predominantly $H_2O$ ice particles. Most of the observations of Type 1 PSCs can be subclassified as Type 1b liquid particles or Type 1a solid particles. Other types of particles have been proposed to explain some specific observations.

* It is now generally accepted that Type 1b PSCs are supercooled ternary solution (STS) droplets that form from SSA without a nucleation barrier. Type 1a PSC particles are generally interpreted as solid nitric acid trihydrate (NAT), but understanding of the phase transition mechanisms leading to their formation is still poor. Better understanding of Type 1a PSCs is needed because solid particles play a significant role in denitrification.

* Many of the Type 1b PSC observations occurred during ongoing fast synoptic cooling events, shortly after the air parcels experienced cold temperatures. Type 1a PSCs, in contrast, have been observed when synoptic temperatures were below the NAT existence temperature for several days. It now appears that theoretical models of Type 1a PSC formation may require knowledge of the air parcel thermal history.

* Mesoscale temperature fluctuations, especially over mountain ranges where such fluctuations can reach 20 K peak-to-peak, are important in PSC formation processes, particularly in the Arctic. The integral effect of such phenomena on polar ozone depletion is still unclear.

* Increases in source gases and cooling of the lower stratosphere from ozone depletion and increasing greenhouse gases favor increased formation and persistence of PSCs. However, an upward trend in PSC occurrence is not discernible in the present satellite data record due to the relatively short length of the record as well as the large variability in cloud sightings from year to year.
**3.1 INTRODUCTION**

Since the discovery of the Antarctic ozone hole, a
collection of field observations, laboratory measure-
ments, and modeling studies has provided compelling
evidence that suspended particles play a key role in halo-
gen activation and nitrogen deactivation and, hence,
ozone depletion in the lower stratosphere. Two major
classes of particles are observed in this altitude region:
the stratospheric sulfate aerosol (SSA), which are present
at all latitudes and are enhanced periodically by volcan-
ic eruptions, and polar stratospheric clouds (PSCs), which
form during both polar winters and on into the austral
polar spring. Other less abundant stratospheric particle
types include cirrus clouds, soot, and rocket exhaust par-
ticles.

The effect of these particles on ozone chemistry is
highly dependent on their concentration, size, phase, and
composition. Hence, an understanding of ozone trends
requires an understanding of these characteristics. Many
measurements have been obtained in recent years that
have increased our knowledge of stratospheric particles.
In parallel, various techniques have been developed to
derive from these measurements the particle properties
needed to represent heterogeneous processes in assess-
ment models, primarily surface area and volume. Fur-
thermore, the 1991 eruption of Mt. Pinatubo, one of the
strongest volcanic eruptions of this century, provided the
opportunity to better assess the nature of volcanic SSA,
its dispersal and decay, and its effects on ozone under
conditions of increasing anthropogenic chlorine. This
chapter reviews our knowledge of stratospheric particles
and assesses our ability to evaluate and predict their for-
mation and evolution. It deals with basic thermodynam-
ics of the major classes of particles, their sources and
possible trends, and describes in situ and remotely sensed
particle observations. Outstanding issues in stratospheric
particles are highlighted as well. Chapter 7 provides
more discussion on particle microphysics and treats the
role of particles in ozone depletion, including chlorine
activation, denitrification, and dehydration. The pos-
sible influence of aircraft emissions on stratospheric
particles will be dealt with in a forthcoming Intergov-
ernmental Panel on Climate Change special report (IPCC,
1999) and, hence, will not be discussed in detail herein.

**3.1.1 Types of Stratospheric Particles**

The most prevalent stratospheric particles are the
sub-micron-sized liquid SSA droplets composed of sul-
furic acid (H₂SO₄) and H₂O. They are present continu-
ously at all latitudes between the tropopause and about
30 km. During volcanically quiescent periods, the ver-
tical distribution of SSA relative to tropopause height is
very similar at all latitudes, with a typical radius (r) of
0.1 µm, mass mixing ratio of 10⁻⁹, and number density
of 10 cm⁻³ (Turco *et al.*, 1982; Hofmann, 1990a).

The other major class of stratospheric particles is PSCs, which are observed in cold regions of the lower polar stratosphere, primarily during winter. On the basis of their optical properties, PSCs have been further divided into two distinct subclasses: Type 1 PSCs that are thought to be relatively small particles containing nitric acid (HNO₃) as a major component, and Type 2 PSCs that are larger, primarily H₂O ice particles. Typical mass mixing ratios for Type 1 and Type 2 PSCs are 10⁻⁸ and 10⁻⁶, respectively (Turco *et al.*, 1989).

Cirrus clouds also occur near the tropopause and
perhaps in the lowest part of the stratosphere in middle
and high latitudes, as shown by an analysis of extinc-
tion data from the Stratospheric Aerosol and Gas Ex-
periment (SAGE) II (Wang *et al.*, 1996). Studies by
Borrmann *et al.* (1996, 1997a) suggest that heteroge-
neous reactions on the surfaces of these clouds can
cause changes in chemical composition akin to those
induced by Type 2 PSCs in the polar stratosphere. A
follow-on study by Solomon *et al.* (1997) suggests that
cirrus can also lead to ozone depletion near the tropo-
pause if they occur with sufficient frequency and spa-
tial extent. The largest effect of cirrus on ozone would
be expected near the midlatitude tropopause of the
Northern Hemisphere in summer.

Recently, some attention has also been focused on
soot, because these particles could act as cloud conden-
sation nuclei and also provide reactive surfaces for het-
erogeneous chemistry (Bekki, 1995; Blake and Kato,
1995; Pueschel, 1996). Maximum soot concentrations
greater than 1 ng m⁻³ are found at northern midlatitudes
at altitudes of about 10 km. Some of the soot particles
are expected to end up embedded in H₂SO₄/H₂O solu-
tion via coagulation with H₂SO₄/H₂O aerosols and, pos-
sibly, condensation of gaseous H₂SO₄.
Model calculations indicate that the impact on the stratospheric aerosol layer of solid particles emitted by rocket and Space Shuttle launches is likely to be very modest (Jones *et al.*, 1995; Jackman *et al.*, 1996).

## 3.1.2 SSA and PSC Thermodynamics

### 3.1.2.1 SSA

Production of SSA droplets begins primarily with the oxidation of gaseous sulfur-bearing compounds such as carbonyl sulfide (OCS) or sulfur dioxide ($SO_2$). The end oxidation product, gaseous $H_2SO_4$, has a very low saturation vapor pressure and therefore takes a particulate form in the lower stratosphere. The exact mechanism of gas-to-particle conversion has not been fully elucidated. New SSA could be formed in situ either by homogeneous nucleation or heterogeneous nucleation on solid materials present in the stratosphere, such as volcanic ash, micrometeorites, aluminum oxide, or soot (Turco *et al.*, 1982; Zolensky *et al.*, 1989; Blake and Kato, 1995; Pueschel, 1996). Traces of foreign materials that have been detected in some SSA (Turco *et al.*, 1982) are not necessarily indicative of heterogeneous nucleation, however, because foreign materials can also be introduced through coagulation.

There is some observational evidence for in situ homogeneous nucleation. For example, after the eruption of Mt. Pinatubo in 1991, increases by up to 2 orders of magnitude in the concentration of condensation nuclei (CN) were observed in the volcanic layers, with 95-98% of the particles composed of $H_2SO_4$ and $H_2O$ according to volatility experiments (Deshler *et al.*, 1992). Electron microscope analysis of these particles revealed that most of the fine particles did not show any evidence of solid or dissolved nucleus material (Sheridan *et al.*, 1992), suggesting that homogeneous nucleation was the main aerosol production mechanism in the Mt. Pinatubo plume. Unusually large increases in CN concentration, likely indicative of nucleation events, have also been detected at about 25-30 km during winter and early spring in northern middle and polar latitudes (Hofmann and Rosen, 1982; Hofmann *et al.*, 1985; Hofmann, 1990b), as well as during spring over Antarctica (Hofmann *et al.*, 1988; Wilson *et al.*, 1989; Deshler *et al.*, 1994a). Such increases in CN concentration seemed to be more intense during volcanic periods. Model calculations suggest that these enhanced CN layers are produced by binary homogeneous nucleation triggered by either subsidence of sulfur-rich air in the cold polar vortex (Zhao

*et al.*, 1995) or by a rapid cooling following a polar stratospheric warming (Hofmann and Rosen, 1983; Bekki *et al.*, 1997).

Overall, conditions favorable for homogeneous SSA nucleation, such as highly enhanced gaseous $H_2SO_4$ levels, rapid temperature fluctuations around the evaporation/condensation point, and very low pre-existing particle surface area density, occur infrequently in the stratosphere except in the early stages of a volcanic plume or at the top of the SSA layer. Atmospheric observations and calculations suggest that the upper tropical troposphere is a major source of CN and as such governs the number concentration of SSA in the lower tropical and midlatitude stratosphere (Brock *et al.*, 1995).

Once formed, SSA droplets grow by coagulation and co-condensation of $H_2SO_4$ and $H_2O$. They are removed from the stratosphere through a combination of gravitational sedimentation and stratospheric-tropospheric exchange. As shown by modeling studies (e.g., Turco *et al.*, 1979; Bekki and Pyle, 1992; Pitari *et al.*, 1993; Tie *et al.*, 1994; Weisenstein *et al.*, 1997), the interplay between growth by condensation, coagulation, and removal of the large aerosol particles by sedimentation leads to the establishment of the well-known unimodal lognormal particle size distribution with a typical mode radius of about 0.08 µm and a typical geometric standard deviation of about 1.8. At typical stratospheric conditions ($200 K < T < 240 K$), liquid SSA have theoretical compositions ranging from 60 to 80% $H_2SO_4$ by weight. Because $H_2O$ is present in much higher concentrations in the atmosphere than $H_2SO_4$, SSA composition can be assumed to be determined by temperature and humidity only (Steele and Hamill, 1981; Carslaw *et al.*, 1997).

Frozen $H_2SO_4$ hydrates are the thermodynamically stable phases in the stratosphere (Gable *et al.*, 1950; Zhang *et al.*, 1993). Sulfuric acid tetrahydrate ($SAT = H_2SO_4 \cdot 4H_2O$) is the favored hydrate under polar winter conditions but is expected to melt at about 220 K upon heating (Middlebrook *et al.*, 1993; Zhang *et al.*, 1993). However, most observations suggest that SSA are, indeed, supercooled liquid droplets rather than solid crystals. This is consistent with laboratory experiments showing that, despite the thermodynamic stability of $H_2SO_4$ hydrates, liquid $H_2SO_4/H_2O$ solutions do not freeze readily, even when highly supercooled (Molina *et al.*, 1993; Anthony *et al.*, 1995; Koop *et al.*, 1997) or when certain solid nuclei are present (Koop *et al.*, 1995; Biermann *et al.*, 1996). Theoretical studies reach a simi-
lar conclusion, i.e., that homogeneous freezing of a pure H₂SO₄/H₂O solution to form SAT is hindered energetically (Luo *et al.*, 1994; MacKenzie *et al.*, 1995). Freezing of H₂SO₄/H₂O solutions might be somewhat facilitated by HNO₃ uptake at low temperatures close to the frost point, Tᵢice (Molina *et al.*, 1993; Beyer *et al.*, 1994; Iraci *et al.*, 1995), although other studies suggest that this might not be sufficient for droplets to freeze (Koop *et al.*, 1995). A temperature hysteresis in the life-cycle of SSA particles with freezing close to Tᵢice and melting at about 215-220 K has been indicated from balloonborne backscatter measurements (Larsen *et al.*, 1995).

### 3.1.2.2 TYPE 1 PSCs

According to the phase diagram of Hanson and Mauersberger (1988) and in the absence of phase transition barriers, gaseous HNO₃ and H₂O would be expected to co-condense to form a stable solid phase, nitric acid trihydrate (NAT = HNO₃·3H₂O) at about 195 K in the polar stratosphere. Because homogeneous nucleation of HNO₃ and H₂O vapors is not favored energetically, pre-existing solid particles, possibly SAT, were previously believed to act as condensation nuclei for NAT. However, laboratory and theoretical studies later indicated that heterogeneous nucleation of NAT on SAT is rather unlikely even at high NAT supersaturations (Iraci *et al.*, 1995; MacKenzie *et al.*, 1995). Some pre-activation of

SAT by a previous exposure to NAT may, however, reduce the level of supersaturation required for NAT condensation (Zhang *et al.*, 1996). These results combined with the fact that liquid H₂SO₄/H₂O solutions do not freeze readily suggest that particles may remain liquid well below the NAT existence temperature (T_NAT) and solid particles, composed of acid hydrates and ice, only form at lower temperatures close to or below Tᵢice.

Thermodynamic models using laboratory data predict a steep increase in HNO₃ solubility into liquid H₂SO₄/H₂O solutions a few degrees below T_NAT (Beyer *et al.*, 1994; Carslaw *et al.*, 1994; Tabazadeh *et al.*, 1994; Luo *et al.*, 1995), leading to the formation of supercooled ternary (HNO₃/H₂SO₄/H₂O) solution (STS) droplets. This mechanism is illustrated in Figure 3-1, where the equilibrium state of a droplet ensemble is shown between 260 K and 185 K at the 55-hPa pressure level. Above approximately 260 K, the liquid fully evaporates, which explains the evaporation of droplets at the upper boundary of the Junge particle layer, while at 3-4 K below the frost point, H₂O ice and possibly acid hydrates precipitate. From 250 K to below 188 K, liquid droplets can exist in the stratosphere and undergo some remarkable changes. Below 250 K virtually all the H₂SO₄ condenses, leading to small, highly concentrated H₂SO₄/H₂O droplets that dilute as the temperature decreases (Steele and Hamill, 1981). At 196 K, T_NAT is reached (arrow in Figure 3-1a), and NAT would be expected to grow readily

**Figure 3-1.** Partitioning of stratospheric gases into liquid stratospheric aerosols: (a) total aerosol volume; (b) H₂SO₄ and HNO₃ liquid concentrations in weight percentage; (c) same as (b) but on logarithmic scale and including hydrochloric acid (HCl), hypochlorous acid (HOCl), and hydrogen bromide (HBr) concentrations. Results for T < 210 K are from Carslaw *et al.* (1997), and those for H₂SO₄ at T > 210 K are from Ayers *et al.* (1980). Atmospheric conditions are 55 hPa pressure, 5 ppm H₂O, 10 ppb HNO₃, 0.5 ppb H₂SO₄, 1 ppb HCl, 10 ppt HOCl, and 10 ppt HBr. The frost point (Tᵢice) and NAT existence temperature (T_NAT) are indicated by vertical arrows. In addition to total volume, panel (a) shows ER-2 measurements of 24 January 1989 (Dye *et al.*, 1992) and the modeled aerosol volume in the observation window of the forward scattering spectrometer probe (FSSP) instrument (0.2 µm < r < 10 µm).

<figure><img src="image_1.png" /></figure>
in the absence of a nucleation barrier, accompanied by a
sharp increase in particle volume and an almost com-
plete removal of HNO₃ from the gas phase. In contrast,
the liquid phase, even at 193 K is still mainly H₂SO₄/
H₂O. However, at lower temperatures between 193 K
and 190 K, increasing solubilities enable a strong,
coupled uptake of gas-phase HNO₃ and H₂O. This up-
take transforms the solution within this interval from
quasi-binary H₂SO₄/H₂O into quasi-binary HNO₃/H₂O
with about 40% HNO₃ by weight and less than 3% H₂SO₄
by weight. The liquid volume grows by a factor of 10,
as was observed during the flight of the National Aero-
nautics and Space Administration (NASA) Earth Re-
sources-2 (ER-2) aircraft on 24 January 1989 (Dye et
al., 1992) (data points in Figure 3-1a). As the tempera-
ture drops further toward Tice, the remaining HNO₃ is
readily taken up, and the droplets grow to a total volume
higher than that of NAT in equilibrium. Figure 3-1c also
shows the uptake of other species that are of primary
importance for heterogeneous chemistry. Type 1b clouds
are highly efficient processors of halogen species (see
Chapter 7). Because some field observations of Type 1
PSCs cannot be explained in terms of NAT or STS clouds
in equilibrium, other metastable phases in equilibrium
(Tabazadeh et al., 1995) or highly nonequilibrium par-
ticles (Meilinger et al., 1995; Peter, 1997) have been
proposed as explanations. STS droplets or other meta-
stable particles would not coexist for long with NAT
particles because NAT particles would be expected to
grow slowly at the expense of the other particles due to
lower NAT saturation vapor pressures.

3.1.2.3 TYPE 2 PSCs

It is now fairly well established that Type 2 PSCs
are primarily water ice particles. Laboratory experiments
(Koop *et al.*, 1995) suggest that ice nucleates inside liq-
uid Type 1 PSC particles or on the surface of solid Type
1 PSC particles once the temperature falls to 2-4 K be-
low Tice. The reason for this rather well-constrained re-
lationship between temperature and ice nucleation is the
onset of strong uptake of H₂O by binary or ternary solu-
tions at these temperatures, leading to highly diluted,
water-like droplets that must freeze. A supercooling of
about 4 K below Tice has recently also been observed
(Carslaw *et al.*, 1998) during the formation of
mountain-wave-induced Type 2 clouds (see Sections
3.4.3 and 7.4.1).

Once formed, ice particles develop further accord-
ing to well-known thermodynamics. However, despite

the close coupling between temperature and ice forma-
tion, the remaining uncertainty in supercooling is still a
critical issue. Whether ice formation begins at 2 K or at
4 K of supercooling determines the extent of ice cloud
occurrence, particularly in the Arctic, and is also likely
to influence the particle number density in individual
clouds and, thus, their optical and chemical properties.
This is similar to open issues in our understanding of
cirrus cloud formation in the upper troposphere.

After ice nucleation from liquid ternary solution
droplets, the fate of the acidic components (HNO₃,
H₂SO₄) is not well known. In some particles they might
freeze as NAT and SAT (Koop *et al.*, 1995, 1997), while
in others some of the HNO₃ might partition back into
the gas phase and the remaining acid might stay as a
liquid coating on the ice particle (Carslaw *et al.*, 1998).
Furthermore, it is not clear whether soluble species such
as HNO₃ could be trapped within the ice matrix or to
what extent such species might lead to a solid coating of
ice particles, influencing their microphysical properties
(see Chapter 7).

**3.1.3 Particle Source Gases**

3.1.3.1 SULFUR

The predominant source of SSA in the stratosphere
is strong, sulfur-rich volcanic eruptions, which are by
nature highly intermittent and unpredictable. The aver-
age flux of volcanic sulfur (S) to the stratosphere over
the last 200 years has been estimated to be around 1 Tg
yr⁻¹ S, with lower and upper bounds of 0.3 and 3 Tg yr⁻¹
(Pyle et al., 1996). A minimum flux of 0.5-1 Tg yr⁻¹ S
for the past 9000 years has been derived from ice core
sulfate data. There have been two large sulfur-rich erup-
tions in the last two decades: El Chichón in 1982 (3.5 Tg
S) and Mt. Pinatubo in 1991 (9 Tg S).

OCS oxidation is believed to be the major non-
volcanic source of stratospheric sulfur (Crutzen, 1976).
Recent estimates of this source range from 0.03 Tg yr⁻¹
S (Chin and Davis, 1995) to 0.049 Tg yr⁻¹ S (Weisenstein
et al., 1997). Although most OCS sources are natural,
there are some indications that anthropogenic emissions
of OCS may be substantial, and possibly increasing
(Khalil and Rasmussen, 1984; Zander et al., 1988;
Hofmann, 1990a). However, historical data on indus-
trial releases suggest that anthropogenic emissions of
OCS and its precursor carbon disulfide (CS₂) have been
relatively constant over the 1977-1992 period (Chin and
Davis, 1993). Furthermore, no statistically significant
trend in lower stratospheric OCS was inferred from
spaceborne observations made in 1985 and 1994
(Rinsland *et al.*, 1996).

Other surface sources of sulfur are approximately 3 orders of magnitude larger than the OCS source. Although most reduced sulfur gases other than OCS are rapidly oxidized to SO₂ and therefore are not expected to reach the stratosphere in large amounts, 2-D model calculations suggest that the stratospheric aerosol loading may be strongly influenced by upper tropospheric sulfur (Weisenstein *et al.*, 1997). However, large uncertainties are associated with tropospheric sulfur chemistry, tropospheric removal, vertical transport of short-lived species (Langner and Rodhe, 1991), and stratospheric-tropospheric exchange (Holton *et al.*, 1995). Such uncertainties reduce the level of confidence in model assessments of the impact of the massive surface emissions of short-lived sulfur (CS₂, dimethyl sulfide (DMS), hydrogen sulfide (H₂S), SO₂) on the stratosphere.

3.1.3.2 NITRIC ACID

Nitrous oxide (N₂O) is the principal precursor of stratospheric reactive nitrogen (NO_y) and hence of HNO₃. Anthropogenic emissions of N₂O stem primarily from agricultural sources and are about half as large as natural emissions. The rate of NO_y production from N₂O oxidation is approximately 0.5-1 Tg yr⁻¹ N (Kasibhatla *et al.*, 1991; WMO, 1995; Vitt and Jackman, 1996). N₂O concentrations are increasing at a rate of about 0.25% yr⁻¹ (WMO, 1995; IPCC, 1996). Lightning may also be another significant, though uncertain, source of nitrogen in the tropical lower stratosphere (Kotamarthi *et al.*, 1994).

3.1.3.3 WATER VAPOR

The main sources of H₂O to the stratosphere are thought to be injection through the tropical tropopause and in situ production from methane (CH₄) oxidation (Harries et al., 1996). Stratospheric-tropospheric exchange at middle and high latitudes might also be important. Direct evidence of the link between the tropical H₂O distribution and the annual cycle in tropopause temperature has been provided by satellite data (Mote et al., 1995, 1996). The average H₂O mixing ratio of air entering the stratosphere inferred from satellite, aircraft, and balloon data ranges from 3.2 to 4.2 parts per million (ppm) (Dessler et al., 1994; Abbas et al., 1996; Engel et al., 1996; Mote et al., 1996; Remsberg et al., 1996).

Because, to first order, two molecules of H₂O can be
formed from one of CH₄, the quantity H₂O + 2CH₄ is
approximately conserved over the lower and middle
stratosphere at a value of 6.7 to 7.6 ppm, as shown by
satellite and aircraft observations. Tropospheric CH₄ has
been increasing at an average rate of 0.6% yr⁻¹ over the
last 10 years (IPCC, 1996). Because production of H₂O
by CH₄ oxidation is relatively small in the lower strato-
sphere, any H₂O trend there should be substantially
smaller than the CH₄ trend.

A statistically significant upward trend in H₂O con-
centration has been detected at altitudes from 16 to 26
km at northern midlatitudes from balloonborne hygrom-
eter measurements carried out from 1981 to 1994
(Oltmans and Hofmann, 1995). The maximum trend was
found to be 0.8% yr⁻¹ at 18-20 km. This finding is cor-
roborated by an apparent increase in H₂O + 2 CH₄ de-
rived from a variety of measurements made since 1975
(Engel et al., 1996). However, the H₂O trend seems to
be larger than the trend expected from the oxidation of
increased CH₄ alone (Oltmans and Hofmann, 1995;
Engel et al., 1996). Nedoluha et al. (1998) also reported
an upward trend in H₂O in the stratosphere based on
Halogen Occultation Experiment (HALOE) and ground-
based millimeter-wave spectrometer data.

3.2 INSTRUMENTS

3.2.1 Measurement Principles

Instruments to measure SSA can be broadly divided into two categories, those that provide ensemble measurements and those that provide individual particle measurements. Ensemble measurements detect the signature of a population of particles at one instant and include passive extinction measurements, such as from the sunphotometer (Volz and Goody, 1962), and active scattering measurements, such as lidar (Fiocco and Grams, 1964). Individual measurements include those that use optical scattering to count and size single particles (Rosen, 1964; Baumgardner et al., 1992) and those that use impaction to collect particles that can then be counted, sized, and chemically analyzed (Bigg, 1975; Farlow et al., 1979). The range of instruments available allows for a variety of measurement platforms. Ensemble measurements are usually conducted on the ground for high temporal and vertical resolution, or from satellites for global coverage. Single-particle measurements are usually conducted using aircraft for high vertical and
horizontal resolution, or balloons for high vertical reso-
lution and accessibility to regions of the stratosphere
above aircraft altitudes.

3.2.2 Contemporary Data Records

Regular measurements using ground-based lidar began in the early 1970s, and some of these records extend to the present (Jäger, 1991; Osborn et al., 1995; Barnes and Hofmann, 1997). Lidars provide high-vertical-resolution measurements of particle backscatter, and often depolarization ratio as well, at one or more wavelengths. Lidar sites now range in latitude from 90°S to 79°N, with a number of sites in the low to midlatitudes, and a few stations in the subtropics. Sunphotometer measurements began in the 1950s and extend to the present. Aerosol optical depth is usually measured at several wavelengths and can then be used to infer column size distributions. Long-term ground-based measurements are available from polar (Herber et al., 1993) and tropical regions (Russell et al., 1993a), and sunphotometers have been deployed on aircraft as well (Russell et al., 1993b).

Aerosol extinction measurements from spaceborne platforms began in 1975 (Pepin et al., 1977). These became routine with the launch of the autonomous limb-viewing optical extinction instruments, Stratospheric Aerosol Measurement (SAM) II and SAGE, in 1979 (McCormick et al., 1979). The near-global, multi-wavelength measurements from SAGE and its successor SAGE II cover the periods 1979-1981 and 1984 to the present, respectively. The single-wavelength record from SAM II, whose spatial coverage was limited to high latitudes, extends continuously from 1978 to 1991, with intermittent data thereafter until late 1993. Satellite extinction measurements were expanded in 1991 with the launch of the Upper Atmosphere Research Satellite (UARS), carrying HALOE, which measures infrared limb extinction, and CLAES (Cryogenic Limb Array Etalon Spectrometer) and ISAMS (Improved Stratospheric and Mesospheric Sounder), which measure infrared limb emission (see Geophysical Research Letters, Vol. 20, No. 12, 1993). The lifetimes of CLAES and ISAMS were limited to less than two years, but HALOE is still operational. High-resolution middle-infrared measurements of extinction by Mt. Pinatubo SSA were made by the Atmospheric Trace Molecule Spectroscopy (ATMOS) instrument during a March-April 1992 Space Shuttle mission (Rinsland et al., 1994). High-latitude

solar occultation aerosol measurements were also made
from 1993-1996 by the Polar Ozone and Aerosol Meas-
urement (POAM) II instrument on the French SPOT (Sat-
ellite Pour l'Observation de la Terre) 3 satellite (Glaccum
et al., 1996). More recently there are new infrared ex-
tinction measurements available from ILAS (Improved
Limb Atmospheric Spectrometer) on the Japanese
ADEOS (Advanced Earth Observing Satellite) platform
from August 1996 to June 1997. The nadir-
viewing AVHRR (Advanced Very High Resolution Ra-
diometer) satellite instruments measure the reflected solar
radiance and infer aerosol optical thickness (Griggs,
1983). These observations have a fine horizontal reso-
lution over oceans for tracking dense volcanic clouds
but cannot provide any vertical resolution (Stowe et al.,
1992).

Individual particle measurements are obtained in situ with aerosol spectrometers. Spectrometers deployed on aircraft include ones with sampling volumes external to the instrument (0.15 µm < *r* < 10 µm) (Baumgardner et al., 1992) and those with internal sampling volumes (0.03 µm < *r* < 1.0 µm) (Jonsson et al., 1995). Particle counters carried on balloons require internal sampling volumes (0.15 µm < *r* < 10 µm) (Hofmann et al., 1975; Hofmann and Deshler, 1991). In situ sampling is also done using impaction devices on aircraft (Pueschel et al., 1989; Sheridan et al., 1994) and on balloons (Sheridan et al., 1992). The physical samples collected by the impaction devices are returned to the laboratory for analysis. Regular stratospheric aerosol measurements using aerosol counters carried on balloon platforms have continued since the 1970s (Hofmann, 1990a; Deshler et al., 1997). Aerosol impactors carried on balloons are employed much less frequently (Bigg, 1975; Sheridan et al., 1992). Stratospheric aerosol measurements using aircraft-borne aerosol spectrometers began with the investigations of polar ozone loss in 1987 and 1989 (Dye et al., 1990a; Wilson et al., 1992) and have continued through the eruption of Mt. Pinatubo (Jonsson et al., 1996). Aerosol impactors on aircraft have been used since the late 1970s (Farlow et al., 1979; Goodman et al., 1994). In number concentration, a large fraction of aerosol has sizes below the minimum sizes detectable with optical instruments. To measure the concentration of these CN, diffusion chambers are placed ahead of optical counters to grow all available particles to optically detectable sizes prior to counting (Rosen and Hofmann, 1977; Wilson et al., 1983). Although CN counters do not discriminate particles according to size,
laboratory tests indicate they are sensitive to particles with $r > 0.01$ µm at stratospheric pressures (Wilson *et al.*, 1983). The first regular sampling of stratospheric CN began in the mid-1970s from balloons and in the 1980s from aircraft. A balloonborne backscattersonde combines in situ methods with an ensemble measurement (Rosen and Kjome, 1991). These instruments have been used in the 1990s for regular midlatitude measurements (Rosen *et al.*, 1994a) and for seasonal Arctic measurements (Larsen *et al.*, 1997).

### 3.2.3 Derivation of Particle Characteristics

For many applications the directly measured particle properties, i.e., extinction, scattering, number, or size, are not the quantities required. In particular, heterogeneous chemical reaction rates may be a function of aerosol surface area or volume, depending on the ratio of the molecular diffusion coefficient to the molecular loss rate coefficient (Hanson *et al.*, 1994). For example, the reaction of chlorine nitrate ($ClONO_2$) with hydrochloric acid (HCl) dissolved in SSA seems to depend on aerosol volume, whereas the hydrolysis of dinitrogen pentoxide ($N_2O_5$) on SSA is fast enough that only aerosol surface area is important. At cold temperatures, particle phase is also important, because the effective uptake coefficient is smaller on solid hydrates than on liquid particles at the same temperature (Ravishankara and Hanson, 1996; Borrmann *et al.*, 1997b). Phase is also important when considering particle growth rates, because STS droplets can be quite a bit smaller than solid $HNO_3$ hydrate particles.

#### 3.2.3.1 SURFACE AREA AND VOLUME

Aerosol surface area and volume concentrations can be obtained from individual particle measurements either by directly integrating the measured discrete size distribution (Dye *et al.*, 1992; Pueschel *et al.*, 1992; Wilson *et al.*, 1993; Goodman *et al.*, 1994; Jonsson *et al.*, 1995) or by fitting measured particle number concentrations with an analytic size distribution function and integrating (Deshler *et al.*, 1993). The uncertainties of 30-50% estimated for surface area and volume obtained in this way arise from Poisson counting statistics at low concentrations, and from sizing errors. The sizing errors (10-20%) arise primarily from uncertainties in particle refractive index, from inhomogeneities in particle illumination, and from spectral broadening (Baumgardner *et al.*, 1992).

Discrete size distributions measured simultaneously by several individual particle measurement techniques have been compared and found to agree favorably when differences in sampling technique are accounted for (Pueschel *et al.*, 1992).

Because the wavelength dependence of extinction is related to particle size (Ångström, 1908), in situ individual and remote ensemble particle measurements can be compared using Mie scattering theory. One approach is to compare satellite extinction measurements to extinctions calculated using discrete and fitted size distributions derived from particle counter data (Wilson *et al.*, 1992; Jonsson *et al.*, 1995; Hervig *et al.*, 1996). The opposite approach is to derive particle size information by inverting multi-wavelength extinction, scattering, or emission data. Twomey (1974) and Capps *et al.* (1982) concluded that extinction measurements at a few wavelengths cannot be expected to yield detailed information about the particle size distribution, but can be used to reliably estimate moments of the distribution. The inversion of multi-wavelength extinction/scattering measurements is complex, and a number of techniques have been tried (Grainger *et al.*, 1995; Yue *et al.*, 1995; Anderson and Saxena, 1996). For estimating surface area or volume using four-wavelength SAGE II extinction data, Thomason and Poole (1993) found that principal component analysis (Twomey, 1977) worked well. Uncertainties in volume and surface area estimated in this way were shown to be ~30%. Other recent estimates of surface area and volume have been made using multi-wavelength lidar data (e.g., Del Guasta *et al.*, 1994; Stein *et al.*, 1994; Gobbi, 1995; Wandinger *et al.*, 1995) and HALOE infrared extinction measurements (Hervig *et al.*, 1997). These estimates were made using empirical relationships established from Mie scattering calculations performed with a large set of observed size distributions. Uncertainties on these surface area estimates are around 30% as well.

Figure 3-2 shows comparisons of surface area and volume estimated from in situ optical particle counter data over Laramie, Wyoming (41°N), and from relevant SAGE II and HALOE satellite extinction measurements (Thomason *et al.*, 1997a; Hervig *et al.*, 1998). These comparisons show differences of roughly 50%, with the satellite estimates typically greater than the particle counter data. Another study of closure among various measurement approaches (Russell *et al.*, 1996) showed that during the 2 years following the Mt. Pinatubo eruption, the area-weighted, or effective, particle radius de-
<figure><img src="image_2.png" /></figure>

**Figure 3-2.** Volumes and surface areas inferred from satellite measurements of visible (SAGE II) and infrared (HALOE) extinction compared with balloonborne optical particle counter measurements over Laramie, Wyoming (41°N). The SAGE data are monthly medians over the 35°-45°N latitude band, while the HALOE data are chosen within a time and distance window as described in Hervig *et al.* (1997). A one-to-one agreement line and error bars of 30% are included for reference.

termed by sunphotometers, satellite instruments, air-borne wire impactors, and balloonborne particle counters agreed to within roughly 25%.

### 3.2.3.2 PHASE

Information on particle phase has been obtained to date primarily from active optical scattering measurements using polarized light. Lidar is the most common instrument with this capability, although some backscattersondes have this capability as well. Linearly polarized light is transmitted, and the returned signal is measured in both the parallel-polarized and cross-polarized channels. Depolarization from an ensemble of scatterers in excess of that expected from the molecular atmosphere is an indication that non-spherical particles are present. The standard value for molecular depolarization is 1.4% (Young, 1980), although some systems with narrowband interference filters use a limit of 0.5% (Shibata et al., 1997). Depolarizing aerosol may consist of a mixture of solid and liquid particles, or of solid particles with a variety of forms. Because of size-dependent polarization properties for fixed particle shape (Mischenko and Sassen, 1998), it is difficult to quantify

depolarization as a measure of the non-sphericity of the scattering particles, although some attempts have been made (Flesia et al., 1994; Carslaw et al., 1998). Browell et al. (1990) were the first to report that PSCs above the frost point could be separated into two classes based on depolarization. Stefanutti et al. (1991) and Adriani et al. (1995) made similar observations in the Antarctic. Observations of depolarized signals from Mt. Pinatubo volcanic aerosol were made in the Antarctic (Gobbi and Adriani, 1993), in northern midlatitudes (Vaughan et al., 1994), and in the Arctic (Toon et al., 1993; Godin et al., 1994; Rosen et al., 1994b). The measurements by Toon et al. were above the Mt. Pinatubo ash layer and those by Rosen et al. were obtained in January 1993 after most of the volcanic ash should have been removed from the stratosphere. These observations might therefore indicate the presence of solid H₂SO₄/H₂O particles (see Section 3.4 and Chapter 7).

### 3.2.3.3 COMPOSITION

The composition of SSA was first assessed using impactors to physically collect the aerosol, which could then be sized and chemically analyzed (Junge et al., 1961;
Bigg, 1975). Other early methods used heaters positioned ahead of aerosol counters to evaporate the aerosol (Rosen, 1971). Determining the temperature at which the majority of aerosol evaporated to sizes below the counter limits allowed the H₂SO₄-H₂O composition to be estimated (Hofmann and Rosen, 1983). Both of these techniques have been used for recent measurements of Mt. Pinatubo aerosol (Sheridan et al., 1992; Deshler et al., 1993; Goodman et al., 1994; Sheridan et al., 1994) and of aerosol in the polar regions (Pueschel et al., 1989; Goodman et al., 1997), and to determine the volatile fraction of tropical CN (Brock et al., 1995) and CN in aircraft plumes (Hofmann and Rosen, 1978; Fahey et al., 1995). An alternate technique using ion mass spectrometry to analyze vaporized particles in situ and comparing the resultant H₂SO₄ concentrations with correlative in situ aerosol concentrations has been recently applied (Arnold et al., 1998). Measurements at temperatures above 200 K are all consistent and indicate that SSA are predominantly droplets of highly concentrated H₂SO₄. Analyses of spaceborne infrared extinction and limb emission data on aerosols also show consistency with a concentrated H₂SO₄ composition (Rinsland et al., 1994; Massie et al., 1996).

As mentioned in Sections 3.1.2.2 and 3.1.2.3, there are still uncertainties with regard to the exact composition of PSCs. Recent approaches to determining the composition of these particles include the use of multi-angle aerosol scattering devices (Baumgardner et al., 1996) and comparison of nearly coincident lidar and in situ measurements (Adriani et al., 1995). These methods have both been shown to be consistent when applied to the well-known SSA, and some initial results are available for Antarctic PSC measurements.

## 3.3 SSA OBSERVATIONS

Direct measurements of SSA were first made in 1957 by Junge *et al.* (1961) using balloonborne impactors. The SSA layer is often called the Junge layer in recognition of these measurements. However, the existence of the layer was suggested some 50 years earlier from twilight observations (Gruner and Kleinert, 1927). As discussed previously, systematic measurements of SSA have been made from a variety of platforms since the early 1970s.

### 3.3.1 Volcanic Aerosol

All the contemporary data records clearly illustrate the strong perturbations to SSA levels caused by sulfur-rich volcanic eruptions (Hofmann, 1990a; Jäger, 1991; Chazette et al., 1995; Osborn et al., 1995; Uchino et al., 1995; Thomason et al., 1997b). Sulfur gases in the eruption plumes are oxidized to H₂SO₄, which then condenses into SSA within about 2 months of an eruption. Historical records spanning the last 100 years or more suggest that the most recent 30-year period can be characterized as a relatively active volcanic period. Primarily through the use of historical pyrheliometric data, Stothers (1996) showed that eight major eruptions occurred during the past century. Four of these occurred between 1880 and 1910 (Krakatau, an unidentified eruption, Santa Maria, and Katmai) and four occurred since 1960 (Agung, Fernandina, El Chichón, and Mt. Pinatubo). Between 1910 and 1960, the stratosphere was almost undisturbed by volcanic activity. Junge's measurements of the SSA layer came at the end of this period, but the contemporary long-term records did not begin until after the start of the current volcanically active period. In terms of impact on the stratosphere, the two largest eruptions in the last 100 years were Krakatau first, and then Mt. Pinatubo (Stothers, 1996). Analyses of ice cores from Greenland permit an even longer view (Hammer et al., 1980), which indicates that global volcanism has an approximate 80-year periodicity. The mid-twentieth century lull in volcanism may be the most recent manifestation of this periodicity.

#### 3.3.1.1 DISPERSAL OF VOLCANIC AEROSOL

The two most recent major eruptions have occurred in the tropics: El Chichón (17°N, April 1982) and Mt. Pinatubo (15°N, June 1991). However, since they occurred in different seasons and at different phases in the vertical structure of the quasi-biennial oscillation (QBO), the dispersal and decay of aerosol from the eruptions were different. A notable feature of low-latitude eruptions is the accumulation of aerosol in a tropical stratospheric reservoir (Trepte and Hitchman, 1992). This reservoir has an abrupt, narrow boundary on the winter hemisphere side and a broad boundary on the summer hemisphere side, as noted following the Mt. Pinatubo eruption (Grant et al., 1996; Lambert et al., 1997). Above about 20 km, detrainment of material from the tropical reservoir is related to planetary wave activity. Thus, aerosol is preferentially transported into the winter hemi-
sphere. When QBO easterly winds overlie the equator, the aerosol reservoir remains relatively isolated from midlatitudes because the penetration of planetary waves into the tropics is inhibited, and removal of aerosols from the tropics occurs on a smaller scale along the periphery of the reservoir and just above the tropopause. Such was the case immediately following eruption of both El Chichón and Mt. Pinatubo (Trepte et al., 1993). When QBO westerly winds are prevalent over the tropics, planetary waves can propagate deeper into the tropics and more easily transport aerosol poleward from the tropical reservoir. A large increase in aerosol loading from the El Chichón eruption was observed in the Northern Hemisphere during the autumn of 1982 after the equatorial stratosphere had made the transition to the westerly phase of the QBO (Pollack et al., 1983).

The secondary circulation in the tropics associated with the QBO also has a strong influence on the dispersal of aerosols to higher latitudes (e.g., Plumb and Bell, 1982). In the westerly phase, there is equatorward flow above the shear layer, which leads to descent and a poleward spreading of conserved tracers below it. The descent of the QBO westerly shear in mid-1982 probably enhanced the transport of El Chichón aerosols into the Northern Hemisphere later that year (Hitchman et al., 1994). The secondary circulation is reversed during the easterly QBO phase, so that aerosols tend to be lofted over the equator, as manifested by enhanced upward transport of Mt. Pinatubo aerosol over the equator during the last half of 1991.

Even though both El Chichón and Mt. Pinatubo are at approximately the same latitude, there was a distinct difference within the tropics in the movement of aerosols from these eruptions. In the case of El Chichón, most aerosols remained in the Northern Hemisphere (Pollack et al., 1983; Stothers, 1996) whereas for Mt. Pinatubo, aerosols were rapidly transported south of the equator (Rosen et al., 1994a; Godin et al., 1996; Deshler et al., 1997). Mt. Pinatubo simulations by Young et al. (1994) and Fairlie (1995) showed that meridional circulation resulting from aerosol-induced local heating caused the initial southward movement of SSA across the equator. However, significant aerosol heating was also noted after the eruption of El Chichón, even though no large-scale cross-equatorial drift of SSA was observed. Trepte et al. (1993) noted that strong QBO easterly winds above about 21 km over the equator might have inhibited transport into the Southern Hemisphere immediately after the eruption of El Chichón. In contrast, Mt.

Pinatubo erupted when the easterly phase of the QBO was just descending through 25 km, and the absence of strong easterlies may have allowed significant amounts of material to be transported across the equator at lower altitudes.

A history of the dispersal of Mt. Pinatubo aerosol as measured by Northern Hemisphere ground-based lidars is shown in Figure 3-3 for the first 2 years following the eruption. The progression of SSA northward can be followed, and its intermittent nature in the first 6 months is obvious. The conservative nature of the volcanic aerosol between 6 and 9 months following eruption of Mt. Pinatubo has been illustrated by Borrmann et al. (1995), indicating that microphysical and vertical mixing processes were nearly completed by that time. This is apparent in the lidar data in Figure 3-3. The top of the aerosol layer is near 25 km, with the peak loading in the 16-22 km region. The altitude of the layer top and that of peak loading both decrease toward higher latitudes. Wintertime subsidence is clearly apparent in the northernmost measurements. A global view from SAGE II of the evolution of aerosol surface area from Mt. Pinatubo at two altitudes is shown in Figure 3-4. The relatively quiescent 6 years between eruption of El Chichón and Mt. Pinatubo are apparent, and within this period at 20 km the eruptions of Ruiz (in 1985) and Kelut (in 1990) can be seen. Ruiz erupted in the middle of the westerly phase of the QBO, which is favorable for the spread of material poleward in both hemispheres. Kelut erupted when strong easterlies were present at 20 km, and thus aerosol would be inhibited from crossing the equatorial region, leading to a preferential dispersal of the aerosol southward (Hitchman et al., 1994). At 25 km the tropical stratospheric reservoir (20°S-20°N) is clearly evident (Trepte and Hitchman, 1992). The strength of the tropical barrier is illustrated by the confinement of the highest surface areas in the tropics even though Mt. Pinatubo was at the edge of the northern boundary. The preferential transport of aerosol to the winter hemisphere is also evident in Figure 3-4. The data indicate some hemispheric differences in the midlatitudes after 1994, as do balloonborne measurements (Deshler et al., 1997).

### 3.3.1.2 DECAY OF VOLCANIC AEROSOL

The decay of stratospheric volcanic aerosol proceeds at fairly well-defined rates. An exponential ($e^{-t}$) decay time of $1 \pm 0.2$ years generally characterizes the
<figure><img src="image_3.png" /></figure>

<figure><img src="image_4.png" /></figure>

<figure><img src="image_5.png" /></figure>

July 91

Jan 92

July 92

Jan 93

July 93

**Figure 3-3.** Time-height profiles of backscatter ratio from lidar measurements at Ny-Ålesund (79°N), Observatoire de Haute Provence, France (44°N), and Naha, Japan (26°N).
<figure><img src="image_6.png" /></figure>

**Figure 3-4.** SAGE II aerosol surface area density as a function of time and latitude at altitudes of (a) 20 and (b) 25 km. White regions indicate no data are available. The latitude and time of significant volcanic events are noted. (Adapted from Thomason *et al.*, 1997a.)

behavior of any of several aerosol measurements during the first 3 years following the Mt. Pinatubo eruption: peak backscatter, peak mass, column backscatter, and column mass (Rosen *et al.*, 1994a; Jäger *et al.*, 1996; Barnes and Hofmann, 1997; Deshler *et al.*, 1997). The decay of the peak parameters is smoother than the decay of columnar quantities because the latter are more influenced by air mass transport, fluctuations in tropopause height, and stratospheric-tropospheric exchange (Rosen

*et al.*, 1994a). The peak parameters reflect primarily sedimentation, are much less influenced by global- and synoptic-scale circulation after the initial mixing period, and are not influenced by fluctuations in tropopause height. Column backscatter measurements by Kent and Hansen (1998) show an $e^{-1}$ time from September 1994 to December 1997 that was more than double (1.8 years) that observed during the immediate post-Pinatubo period (0.8 year). The decay of aerosol surface area at three altitudes in the northern midlatitudes following eruption of Mt. Pinatubo is shown in Figure 3-5. The nature of the decay curves and the time required to reach pre-eruption conditions are different at the three altitudes. At 25 km, there is a clear QBO signature, and average surface areas reached pre-Pinatubo values by early 1995, some 3.5 years after the eruption. The decay at 20 km is considerably smoother, but it took an additional year (until early 1996) to reach pre-eruption conditions. The record at 15 km is quite different, clearly showing seasonal changes, namely, increases in winter/spring and decreases in summer/fall. At this altitude, pre-Pinatubo surface areas were not reached until mid- to late 1996, more than 5 years after the eruption. The decay of aerosol surface area is on the order of 20-30% slower than the decay of backscatter or mass (Rosen *et al.*, 1994a; Jonsson *et al.*, 1996; Deshler *et al.*, 1997). Chazette *et al.* (1995) found the decay of integrated backscatter between 15 and 20 km to be on the order of 25% longer after Mt. Pinatubo than after El Chichón eruption. They attributed this slower decay to the higher lofting of particles following eruption of Mt. Pinatubo. Russell *et al.* (1996) showed that the stratospheric aerosol effective radius increased from 0.15 µm prior to the Mt. Pinatubo eruption, to a maximum near 0.55 µm one year later, then decreased to 0.45 µm by spring 1993. The effective radius continued to decrease to a value of <0.2 µm by fall 1994 in the midlatitudes of both hemispheres and has remained between 0.15 and 0.2 µm since then (Deshler *et al.*, 1997).

Number size distributions at 20 km for the first five springs following Mt. Pinatubo eruption are shown in Figure 3-6 and are compared with a pre-Pinatubo measurement, when the data were well represented by a unimodal lognormal distribution. After the eruption the data are represented with a bimodal lognormal distribution. The two modes of the distribution were initially quite narrow, and the median radius in the small-particle mode ($r_1$) increased by a factor of 2. In the next 3 years $r_1$ decreased, and the width ($\sigma_1$) in the first mode in-
<figure><img src="image_7.png" /></figure>

**Figure 3-5.** History of the evolution of aerosol surface area at 25, 20, and 15 km following the Mt. Pinatubo eruption. Surface areas estimated from balloonborne in situ optical particle counter (OPC) measurements above Laramie (41°N) and from SAGE II data in the 35°-45°N latitude band are shown. The error bars on the in situ data represent an uncertainty of 30%. The SAGE II error bars represent the standard deviation about the zonal median. (Adapted from Thomason *et al.*, 1997a.)

creased. The particle concentrations ($N_1$, $N_2$) and surface areas ($A_1$, $A_2$) shown for the springs of 1992 and 1993 compare favorably with ER-2 measurements by Jonsson *et al.* (1996). The second (large-particle) mode is well defined throughout the 5 years shown in Figure 3-6, preserving a quite similar shape. The primary change in the second mode is in the number concentration ($N_2$), which decreased steadily by about a factor of 5 per year. The second mode had almost disappeared by spring of 1997.

Quite similar decay rates have been observed for SSA from the El Chichón and Mt. Pinatubo eruptions (Rosen *et al.*, 1994a; Chazette *et al.*, 1995; Barnes and Hofmann, 1997), and from the decay at different latitudes for Mt. Pinatubo aerosol (Jayaraman *et al.*, 1995; Jäger *et al.*, 1996). These decay rates are also similar to that observed for the removal of strontium from the stratosphere (Fabian *et al.*, 1968). These similarities reflect the fact that the removal of SSA, once it is dispersed meridionally, is controlled by relatively steady and robust processes, such as gravitational settling and stratospheric-tropospheric exchange. The maximum stratospheric-tropospheric exchange of tracers is in the

Northern Hemisphere spring (Holton *et al.*, 1995), which may explain the clear signature of a surface area minimum at 15 km in Figure 3-5 during that season each year. The temporal and vertical resolution required to capture stratospheric-tropospheric exchange processes in detail is provided by lidar, and several such studies have pointed out the influence of volcanic SSA on the upper troposphere (Menzies and tratt, 1995; Sassen *et al.*, 1995; Post *et al.*, 1996).

### 3.3.2 Background Aerosol

Since the discovery of the SSA layer in 1957 (Junge *et al.*, 1961) there has been much speculation about the stability of the layer and the source of the $H_2SO_4$ that is the primary component of the aerosol. The measurements by Junge *et al.* were made at the end of a lengthy period free of volcanic eruptions (Stothers, 1996), but were not extensive enough to establish a baseline. There are four periods in the modern (post-1970) measurement era when the influence of volcanic eruptions has been at a minimum: 1974, 1979, 1989 to early 1991, and the present (1998). Many studies have focused on these data
<figure><img src="image_8.png" /><figcaption>Figure 3-6. Unimodal/bimodal differential (n) and integrated (N) lognormal size distributions fitted to 1-km averaged particle counter measurements (solid circles) at 20 km above Laramie for the springs of 1991-1996. The parameters of the size distributions and the inferred surface area (A) and mass mixing ratio (m), assuming a composition of 65% H₂SO₄ by weight, are shown on the figure. Uncertainties in the concentration measurements are shown by the vertical bars on the right-hand side. The goodness of the fit is shown at the bottom as the sum of the log of the ratios of measured and calculated number concentration (N_m, N_c).</figcaption></figure>

periods in an attempt to clarify the processes by which the background non-volcanic SSA layer is sustained and to explain the cause(s) for changes observed between the recent minimum periods.

In situ measurements at 41°N presented by Hofmann (1990a) showed an SSA increase of 30-50% between 1979 and 1989, a result corroborated by SAGE data at this latitude (Thomason *et al.*, 1997b). It was suggested that increased surface emissions of OCS or SO₂ (Hofmann, 1990a) or sulfur from aircraft exhaust (Hofmann, 1991) might be responsible for the observed SSA increase. Model results by Bekki and Pyle (1992)

suggest that although aircraft may be a substantial source of sulfate below 20 km, the rise in air traffic from 1979 to 1989 was too small to account for the observed increase in SSA. More recent studies suggest also that the amount of stratospheric sulfur derived from OCS oxidation (Chin and Davis, 1995) is too small to sustain the observed minimum SSA loading. This led Chin and Davis to question whether the minimum SSA periods in the modern data record have been truly free of volcanic effects. Near-global SAGE II data, in fact, show that the 1989-1991 minimum period was influenced by the smaller volcanic eruptions of Nevado del Ruiz in 1985
<figure><img src="image_9.png" /></figure>

**Figure 3-7.** Integrated aerosol backscatter at a wavelength of 0.6943 µm from lidar measurements at Garmisch-Partenkirchen (48°N), NASA Langley (37°N), and Mauna Loa (20°N) for the period 1974-1998. Also shown for the period 1991-1997 is integrated aerosol backscatter calculated from size distributions fitted to in situ data above Laramie (41°N). The calculations assume particles with an index of refraction of 1.45. Also shown are column integrals of surface area from the 1991-1997 in situ data. The column integrals are from the tropopause to near 33 km for all the data.

and Kelut in 1990 (Thomason *et al.*, 1997b).

Figure 3-7 shows column stratospheric backscatter at the ruby wavelength (0.6943 µm) from ground-based lidar sites at Garmisch-Partenkirchen (48°N), NASA Langley (37°N), and Mauna Loa (20°N) for the period of 1974-1998. The Garmisch-Partenkirchen data for 1991-1998 were measured at a wavelength of 0.532 µm and converted to the ruby wavelength using balloonborne particle size information from Laramie. The figure also shows calculated column aerosol backscatter and surface area based on in situ size distribution data for 1991-1997 collected over Laramie. These SSA

records are all in quite good agreement and show that levels in late 1997 were below pre-Pinatubo values, and are likely still decreasing. This means that any anthropogenic contribution must be less than previously estimated based on the observed 1979-1989 increase. Interpretation of the column data at low SSA levels is complicated by relatively large seasonal and QBO-related fluctuations. For example, in late 1996 column backscatter at Mauna Loa dropped as low as any observed in the previous 17 years, but increased in 1997 as the QBO changed phase (Barnes and Hofmann, 1997). Such a modulation is consistent with the tropical upper tropo-
<figure><img src="image_10.png" /><figcaption>**Figure 3-8.** Peak aerosol scattering ratio at a wavelength of 0.6943 µm from lidar measurements at Garmisch-Partenkirchen (48°N) and NASA Langley (37°N) for 1978-1979, 1989-1990, and 1996-1997. Vertical bars are estimated uncertainties in representative measurements.</figcaption></figure>

sphere being a source of CN (Brock *et al.*, 1995) and SO₂ (Weisenstein *et al.*, 1997). The midlatitude column measurements also show this SSA minimum, as well as the increase in 1997 with the QBO returning to a westerly phase. The midlatitude column data have a strong seasonal cycle, i.e., an increase and relative stability during winter and spring followed by a decrease in the summer and fall, that is closely tied to tropopause height. At very low SSA levels, column stratospheric backscatter can fluctuate by factors of 2 to 4 with season.

Long-term records have also been maintained of the peak aerosol scattering ratio, defined as the maximum value in a given lidar data profile of the ratio of aerosol backscatter to molecular backscatter. This quantity fluctuates less with season than do column integrals, and its uncertainty can be estimated quite easily. Figure 3-8 shows peak aerosol scattering ratio data (with representative estimated uncertainty bars) from Garmisch-Partenkirchen and Langley during 1978-1979, 1989-1990, and 1996-1997. Values in late 1997 were clearly

lower than those observed in 1990, but were about 40% higher on average than those observed in early 1979. However, given the uncertainty in the individual measurements and the likelihood that SSA levels are continuing to decrease, the 1979-1997 difference must be viewed with caution at this time. Several more years of data uncorrupted by volcanic activity are required for determining if a stable background has been reached and for a more meaningful comparison with data from the 1979 minimum period.

## 3.4 PSC OBSERVATIONS

Observations of PSCs have been reported for more than a century (Stanford and Davis, 1974). Visual sightings have been mainly in northern Scandinavia during winter, where the clouds occasionally form in wave-like patterns of limited spatial extent when air parcels are lifted and cooled adiabatically over the Norwegian mountains. Due to the nearly monodisperse particle size
and the uniform variation in particle size across these clouds, wavelength-dependent scattering of incident sunlight may cause colorful appearances, giving rise to the name mother-of-pearl, or nacreous, clouds. Systematic satellite measurements later revealed that PSCs can occur on synoptic scales in both the Arctic and Antarctic when stratospheric temperatures fall below about 200 K (McCormick et al., 1982).

### 3.4.1 Distinction between Types of PSCs

Chapter 7 shows that heterogeneous chlorine activation is controlled to first order by temperature and water vapor pressure, and not by detailed particle characteristics. However, in situations where there is not continued exposure to cold temperatures, such as in the Arctic, distinction and understanding of PSC types and details of particle size, composition, and phase are important.

Early ground-based and airborne lidar measurements indicated two distinct growth stages of PSCs (Iwasaka et al., 1985; Poole and McCormick, 1988), giving rise to the classification into Type 1 PSCs, forming 2-6 K above the frost point ($T_{ice}$), and Type 2 PSCs, forming below $T_{ice}$. Lidar depolarization measurements in the Arctic later showed that Type 1 PSCs occur in at least two forms (Browell et al., 1990; Toon et al., 1990): Type 1a PSCs with low backscatter and high depolarization, corresponding to non-spherical, presumably solid particles; and Type 1b PSCs with high backscatter and low depolarization, suggesting spherical, presumably liquid particles (Carslaw et al., 1994; Tabazadeh et al.,

1994). Type 2 PSCs have both high backscatter and high depolarization, corresponding to large ($r \ge 1 \mu m$) ice particles. Subsequent PSC lidar observations in both hemispheres have confirmed these classifications (Stefanutti et al., 1991; Beyerle et al., 1994; Adriani et al., 1995; Stebel et al., 1998; Shibata et al., 1997; David et al., 1998). This is demonstrated in Figure 3-9, which, for comparison, also includes SSA measurements, which exhibit both low backscatter and low depolarization. Table 3-1 summarizes the classification of particles observed in the winter polar stratospheres. Similar classifications of PSC particles have also been obtained from in situ backscatter ratio and color index measurements by balloonborne backscattersondes (Larsen et al., 1997; Rosen et al., 1997).

However, not all winter stratospheric particle observations can be categorized so narrowly. During strong mountain-wave-induced cooling, large particles, presumably Type 2 PSCs, have been observed at temperatures more than 3 K above $T_{ice}$ (Deshler et al., 1994a). Other observations suggest that a nitrate coating could prevent large ice particles from subliming at these elevated temperatures (Gandrud et al., 1990; Wofsy et al., 1990; Goodman et al., 1997). At temperatures above $T_{NAT}$, the ubiquitous SSA is observed by lidar. The observations show large variability in backscatter due to varying aerosol mass loading after volcanic eruptions, but mostly exhibit low depolarization ratios, indicating spherical liquid particles. However, depolarizing and presumably small particles have occasionally been measured by lidars at temperatures above $T_{NAT}$ (Adriani et al., 1995; Stefanutti et al., 1995; Stein et al., 1995; Nagai et al.,

<table><caption>Table 3-1. Classification of stratospheric particle types, based on 0.532-µm lidar backscatter ratios and volume depolarization (David et al., 1998)<fnref n="1" />.</caption><thead><tr><th>Particle type</th><th>Background aerosol</th><th>Volcanic sulfate aerosol</th><th>Type 1a PSC</th><th>Type 1b PSC</th><th>Type 2 PSC</th></tr></thead><tbody><tr><td>Backscatter ratio</td><td>&lt; 1.2</td><td>&gt; 1.2</td><td>&lt; 2</td><td>&gt; 2</td><td>&gt; 2</td></tr><tr><td>Depolarization, %</td><td>&lt;~ 1.5</td><td>&lt;~ 1.5</td><td>&gt;~ 1.5</td><td>&lt;~ 1.5</td><td>10-50</td></tr></tbody></table>

<footnote n="1">¹ The backscatter ratio is defined here as $(B_p+B_m)/B_m$ where $B_p$ and $B_m$ are the particulate and molecular volume backscatter coefficients, respectively. The depolarization ratio is defined as the ratio between the total volume backscatter coefficients in the parallel and perpendicular planes relative to the polarization plane of the emitted laser beam.</footnote>
<figure><img src="image_11.png" /><figcaption>**Figure 3-9.** Scatter plots of 0.532-µm backscatter ratio versus volume depolarization from Arctic measurements at Ny-Ålesund, Spitsbergen, (upper panel) and Antarctic measurements at McMurdo (lower panel), illustrating the grouping of different particle types (classified as in Table 3-1) observed in the polar winter stratosphere. Note the changes of scales on the axes. The Arctic data points are used again in Figure 3-11 to show the temperatures at which the observations were made.</figcaption></figure>
<figure><img src="image_12.png" /></figure>

<figure><img src="image_13.png" /></figure>

<figure><img src="image_14.png" /></figure>

**Figure 3-10.** Example bimodal lognormal size distributions fitted to particle counter measurements (solid circles) in: (a) a non depolarizing PSC; (b) a depolarizing PSC with moderate scattering ratio; and (c) an ice cloud with high depolarization and high scattering ratio. The two differential distributions (long dashes) and the cumulative distribution (solid line) are shown in each panel. The constant unimodal curve (short dashes) in each graph is representative of the size distribution of SSA at the time of these PSC measurements. The parameters of each lognormal size distribution and the inferred surface area and mass are shown on the figures, as are the depolarization (Dep) and scattering ratios (SR) measured by lidar at the time of these in situ measurements. Uncertainties in the concentration measurements are shown by the vertical bars on the right-hand side. Refer to Figure 3-6 and Section 3.3.1.2 for definitions of symbols used in the figure legend.

1997), perhaps indicating solid remnants of previously formed PSC particles.

Because winter temperatures are generally lower in the Antarctic than the Arctic, Type 2 PSCs form much more frequently in the Southern Hemisphere. However, Type 2 PSCs have been observed by lidar and satellite in the Arctic, both under synoptic-scale cooling (McCormick et al., 1990; Pitts et al., 1990) and mountain-wave-induced cooling to below $T_{ice}$ (Müller et al., 1995; David et al., 1997).

### 3.4.2 Measurements of PSC Characteristics

Aerosol depolarization measurements by lidar have provided the best indicator for non-spherical and thereby presumably solid particles in the stratosphere. In order to derive other PSC characteristics from remote sensing data, the refractive index of PSC particles of different composition must be known. Combinations of size distributions measured by optical particle counters and single-wavelength lidar observations have been used to estimate the real part of the refractive index ($m$) for PSC particles (Adriani et al., 1995). At a wavelength of 0.532

$\mu$m, estimated values are $m = 1.42 \pm 0.04$ for NAT, $m = 1.39 \pm 0.03$ for STS (varying directly with temperature as derived from model calculations and multi-wavelength lidar measurements (Luo et al., 1996; Beyerle et al., 1997; Wedekind et al., 1997), and $m = 1.32 \pm 0.01$ for Type 2 PSCs, as expected for ice.

PSC particle size distributions have been measured in situ by several instruments: the airborne forward scattering spectrometer probe (FSSP-300) and multi-angle aerosol spectrometer probe (MASP) (Ferry et al., 1989; Dye et al., 1992, 1996), airborne wire impactors (Goodman et al., 1989, 1997), and balloonborne optical particle counters (Hofmann, 1990c; Hofmann et al., 1990; Rosen et al., 1990; Hofmann and Deshler, 1989, 1990, 1991; Deshler et al., 1991, 1994a, 1994b; Deshler and Oltmans, 1998; Adriani et al., 1992, 1995). Examples of typical PSC size distributions are shown in Figure 3-10. Unimodal Type 1 PSC distributions, with mode radii from 0.3 to 0.5 $\mu$m and representing a large fraction of the available condensation nuclei, are generally observed during faster synoptic cooling (4-10 K day$^{-1}$; Deshler et al., 1991; Beyerle et al., 1994; Adriani et al., 1995). These particles have low depolarization, associated with Type 1b PSCs. In
<figure><img src="image_15.png" /></figure>

**Figure 3-11.** Scatter plots of aerosol backscatter ratio (upper panel) and volume depolarization (lower panel) versus the difference between ambient air temperature $T_{\text{air}}$ and $T_{\text{NAT}}$ from the Arctic lidar observations as shown in Figure 3-9. Type 1b PSCs show a characteristic sharp increase in backscatter ratio at roughly 3-4 K below $T_{\text{NAT}}$, as expected for STS particles (upper panel), whereas Type 1a PSC are observed at all temperatures below $T_{\text{NAT}}$ (lower panel). The transition-type PSC particles are observed at temperatures approaching the ice frost point at roughly 7 K below $T_{\text{NAT}}$. $T_{\text{NAT}}$ was estimated by assuming a high-latitude Northern Hemisphere LIMS (Limb Infrared Monitor of the Stratosphere) $\text{HNO}_3$ profile along with 4.6 ppm $\text{H}_2\text{O}$, and ambient air temperatures were obtained from daily radiosonde measurements.

bimodal Type 1 PSC size distributions, the large-mode radii are around 1.5-2.5 µm, and the large size mode typically represents less than 1% of the available condensation nuclei. These particles are observed during slower synoptic cooling (1-3 K day⁻¹; Beyerle et al., 1994; Adriani et al., 1995), sometimes forming in thin layers, and the particles have large depolarization, associated with Type 1a PSC. Type 2 PSC particles captured in

situ have been predominantly solid and hollow columns and hexagonal plates, with sizes on the order of 5-50 µm and number concentrations of (0.5-5) × 10⁻⁴ cm⁻³ (Goodman et al., 1989).

The first in situ Antarctic PSC measurements were obtained in mid- to late winter 1987; these indicated that the clouds could be composed of NAT (Fahey et al., 1989). Observed size distributions are often bimodal
(Type 1a) (Hofmann and Deshler, 1991; Deshler et al., 1991). Later observations from the Arctic revealed that PSCs often do not form until high supersaturation with respect to NAT is reached, indicating another composition of the particles (Rosen et al., 1989; Schlager et al., 1990). This was particularly the case for PSCs with a unimodal size distribution (Type 1b) where a large fraction of the condensation nucleus population undergoes particle growth (Dye et al., 1990b; Hofmann et al., 1990). From particle volumes derived from size distributions measured with the FSSP, it appeared that SSA could remain in the supercooled liquid state to very low temperatures (down to at least 193 K) and over large regions of the Arctic. The main increase in Type 1 PSC particle volume was observed at temperatures 3-4 K below $T_{\text{NAT}}$ (Dye et al., 1992). In Figure 3-1 measurements of particle volumes from the NASA ER-2 flight on 24 January 1989 were explained in terms of STS solution droplets in equilibrium with the ambient $HNO_3$ and $H_2O$ vapor (Drdla et al., 1994; Tabazadeh et al., 1994; Carslaw et al., 1994, 1995). This explanation is based on thermodynamic models for electrolytic solutions. Ground-based lidar (Beyerle et al., 1997; David et al., 1998), satellite (Taylor et al., 1994; Massie et al., 1997), balloonborne backscatter sonde (Larsen et al., 1997; Rosen et al., 1997), and airborne in situ (Dye et al., 1996; Del Negro et al., 1997) data have provided further evidence that Type 1b PSCs are associated with an STS composition. By relating the lidar measurements of PSC particles to the difference between the air temperature and $T_{\text{NAT}}$, as shown in Figure 3-11, a sharp increase in aerosol backscatter for Type 1b PSCs is observed at temperatures 3-4 K below $T_{\text{NAT}}$, as expected from STS droplets. Also, infrared spectroscopic measurements of PSCs have indicated an STS composition (Toon and Tolbert, 1995), and satellite observations have shown that gas-phase $HNO_3$ is depleted at temperatures where STS droplets could initially be expected to form (Santee et al., 1998). Hence, there is substantial observational evidence that Type 1b PSCs have an STS composition.

Type 1a PSC particles are observed at all temperatures below $T_{\text{NAT}}$ (cf. Figure 3-11), indicating a NAT composition for at least some of the observations (Larsen et al., 1996, 1997; Rosen et al., 1997; David et al., 1998), in agreement with earlier observations from Antarctica (Fahey et al., 1989). Measurements of low column abundances of gas-phase $HNO_3$ by Höpfner et al. (1996) have also indicated the existence of equilibrium NAT particles in the Arctic. The smooth, compact relationship between particle volume and temperature characteristic of STS

droplets (Figure 3-1) is not present in all observations (Dye et al., 1992, 1996; Tabazadeh et al., 1995). Figure 3-12 shows particle counter measurements made on the ER-2 aircraft on 16 January 1989 in the Arctic, similar to those already shown for 24 January 1989 in Figure 3-1a. On the basis of simultaneous in situ measurements of $NO_y$ and particle volume, Tabazadeh and Toon (1996) inferred the composition of a subset of PSC particles observed on 20 January 1989 to be a metastable water-rich $HNO_3/H_2O$ solid phase. They named these particles Type 1c PSCs and suggested them as possible precursors of more stable NAT or nitric acid dihydrate (NAD) Type 1a PSCs. Presently, it is uncertain if the transition-type PSC seen in Figures 3-9 and 3-11 is indicative of Type 1c PSCs. For this same series of ER-2 flights, measurements that deviate from the compact STS curve have also been interpreted as not fully developed NAT clouds since the equilibrium times of NAT particles are relatively long (Peter, 1997). The various panels of Figure 3-12 are explained in detail in Chapter 7, Section 7.2.2.

### 3.4.3 Spatial and Short-Term Variability of PSCs

Many observations suggest that PSCs can exist on synoptic scales, therefore providing a large area through which air parcels can be processed. In addition, meso-scale temperature fluctuations caused by gravity waves or mountain waves can cause PSCs to form on smaller spatial scales.

#### 3.4.3.1 SYNOPTIC-SCALE PSC FORMATION

Synoptic-scale PSCs located outside typical mountain wave regions have been observed by airborne lidar and in situ sensors (e.g., Browell et al., 1990; McCormick et al., 1990; Dye et al., 1992), and by ground-based lidars (e.g., Adriani et al., 1995; Stebel et al., 1998; Shibata et al., 1997; David et al., 1998) and backscattersondes (e.g., Rosen et al., 1993; Larsen et al., 1997).

Tabazadeh et al. (1995) analyzed the temperature histories of a subset of in situ measurements of synoptic-scale PSCs encountered by the ER-2 in January 1989 (Dye et al., 1992). This analysis showed that PSC particles observed on 24 and 25 January that had STS properties (i.e., Type 1b PSCs; cf. Figure 3-1a) had experienced a relatively fast cooling from above the sulfuric acid tetrahydrate (SAT) melting temperature. Other PSC data, from the same series of flights, that deviated
<figure><img src="image_16.png" /></figure>

**Figure 3-12.** Upper large panel: Total particle volume observed by Dye *et al.* (1992) under conditions as in Figure 3-1, but on 16 January 1989. All calculations assume 5 ppm H₂O. Solid line: STS droplets for 10 ppb HNO₃. Dashed lines (top to bottom): STS droplets for 20, 5, 2, and 0 ppb HNO₃. Dotted line: NAT for 10 ppb HNO₃. Lower small panels: Examples of the measured particle size distribution (cases 1-4 refer to the corresponding points in the upper panel). (1) STS droplets at 193.3 K; (2) nonequilibrium particles at 194.8 K (Type 1a); (3) large particles at 190.5 K tentatively identified as NAT-coated ice particles; (4) particles under strongly denitrified conditions at 190.2 K. (Adapted from Peter, 1997.) Insert: Average values and standard deviations of reduced dataset (time window and *r* < 2 µm) interpreted thermodynamically as a new cloud type (Type 1c). (Adapted from Tabazadeh and Toon, 1996.)

from the compact STS curve were not exposed to SAT
melting and had experienced cooling/warming cycles.
Tabazadeh *et al.* (1995) explained this in terms of the
formation of numerous small Type 1c particles prior to
the aircraft encounter (see Section 3.4.2), whereas Peter
(1997) argued that the bimodal character of the particle
distribution would point toward Type 1a particles possi-
bly mixed into a Type 1b cloud. A further analysis of

airborne lidar observations by Tabazadeh *et al.* (1996)
concluded that air masses must be exposed to synoptic
temperatures below *T*<sub>NAT</sub> for 1 day or more in order for
synoptic-scale Type 1a PSCs to form.

Similar analyses of synoptic-scale PSC observa-
tions by ground-based lidar in Antarctica (Adriani *et al.*,
1995) and by backscattersondes in the Arctic (Larsen *et*
al., 1996, 1997) also suggested that most liquid Type 1b
PSC particles are observed during fast synoptic cooling, shortly (<1 day) after entering a cold region. Solid Type 1a PSCs, in contrast, were observed when temperatures had been below $T_{\text{NAT}}$ for a longer time during both synoptic cooling and heating.

### 3.4.3.2 MESOSCALE PSC FORMATION

Gravity waves can cause mesoscale temperature fluctuations of moderate amplitude (~4 K peak-to-peak) that are ubiquitous and must be superimposed on synoptic-scale air parcel trajectories (Murphy and Gary, 1995). Even larger perturbations (>20 K peak-to-peak) with cooling rates greater than 50-100 K h⁻¹ can be encountered in mountain waves (Volkert and Intes, 1992), causing the development of different PSC characteristics on synoptic and regional spatial scales. PSC formation in mountain lee waves has been observed extensively from Kiruna, Sweden, by ground-based and airborne lidars and balloonborne optical particle counters. Strong temperature perturbations, induced by adiabatic lifting of air parcels over the Norwegian mountains and superposed on low synoptic-scale temperatures, clearly influence the cloud formation, as observed by airborne lidars flying parallel to the air flow (e.g., Godin et al., 1994). The temperature may approach or go below $T_{\text{ice}}$ for relatively short periods, causing peculiar PSC features. Particle microlayers with vertical thickness of 75-300 meters have been observed in the Arctic, in which a small fraction (<1%) of the available condensation nuclei have grown into large (2-3 µm) solid particles (Hofmann and Deshler, 1989; Hofmann, 1990c; Stein et al., 1995). Some data in these microlayers are consistent with a NAT particle composition. Other large particles (up to 5 µm) observed by Deshler et al. (1994a) and constituting a larger fraction of the available condensation nuclei must have been composed mainly of ice (taking into account their large size), even though temperatures were well above $T_{\text{ice}}$. Also thin layers of Type 2 PSC, surrounded by thick layers of liquid Type 1b PSC, have been observed (David et al., 1997).

Figure 3-13 gives an impressive quasi-Lagrangian view of PSC formation in lee waves as observed by airborne lidar crossing the mountains near Kiruna in a direction almost parallel to the stratospheric wind. Strong gravity wave perturbations are clearly visible in the particle backscatter (upper panel). These perturbations indicate temperature fluctuations of up to 13 K peak-to-peak with cooling/heating rates of more than 100 K h⁻¹.

Apparently, two distinct and not well understood formation processes of solid particles, identified by large depolarization ratios (lower panel), are captured by these measurements (Carslaw et al., 1998). Between 17°E and 18°E, numerous Type 2 PSC ice particles seem to form directly from liquid Type 1b PSCs. Slightly downstream of this region, liquid PSC particles are again observed. Farther downstream, east of 22°E, a small number of depolarizing (solid) particles with low backscatter ratios (Type 1a PSC) appear to form gradually out of the background aerosol population. This indicates that freezing processes in connection with mountain-wave-induced ice clouds could represent an additional mechanism for Type 1a PSC formation (Meilinger et al., 1995).

## 3.4.4 Seasonal Evolution and Hemispheric Differences

### 3.4.4.1 TEMPERATURE CONDITIONS

In the Arctic, where climatological stratospheric temperatures do not fall below the PSC formation thresholds, the clouds mainly form and disappear repeatedly in connection with transient, synoptic-scale cold outbreaks (Taylor et al., 1994; Pawson et al., 1995) and in mountain wave regions. In addition, individual air parcels often experience temperatures above the SAT melting temperatures during winter, whereby any remnants of solid PSC particle formation would presumably disappear. Therefore, throughout the Arctic winter, temperatures hover around the threshold for PSC formation, and the clouds are generally observed shortly after they initially form, often in liquid or perhaps metastable phases.

Over Antarctica, climatological temperatures are lower than $T_{\text{NAT}}$ for several months and often fall below $T_{\text{ice}}$ as well. PSCs occur continually throughout the winter and, hence, can be observed well after their initial formation period. Individual PSC particles presumably exist for longer periods, which increases the probability of freezing into stable NAT or ice particles. The lower temperatures and longer lifetime of individual PSC particles favor the formation of relatively large particles. This, together with a longer PSC season, leads to more efficient denitrification and dehydration by particle sedimentation, as compared to conditions in the Arctic. Hence, climatologically different temperatures may give rise to hemispheric differences in particle properties and seasonal PSC evolution, and thereby to differences in heterogeneous processing leading to ozone depletion.
<figure><img src="image_17.png" /></figure>

**Figure 3-13.** Altitude/longitude plot of airborne lidar observations from the German Transall aircraft on 15 January 1995 over the Norwegian mountains west of Kiruna, showing 0.532-µm aerosol backscatter ratio (upper panel) and depolarization (lower panel). The wind direction was from west to east (left to right in the figure). (Adapted from Carslaw *et al.*, 1998.)
**Figure 3-14.** Zonally averaged PSC sighting probabilities for the Antarctic from measurements by SAM II for 1979-1989 (upper panel) and POAM II for 1994-1996 (lower panel). (Adapted from Poole and Pitts, 1994; and Fromm *et al.*, 1997.)

<figure><img src="image_18.png" /></figure>

### 3.4.4.2 PSC CLIMATOLOGY

A PSC climatology was developed from more than 10 years of measurements by the Stratospheric Aerosol Measurement (SAM) II satellite instrument (Poole and Pitts, 1994). The frequency of occurrence of Antarctic PSCs observed by CLAES in 1992 generally agreed with this climatology (Mergenthaler *et al.*, 1997). PSC observations in both hemispheres were also made more recently (October 1993-November 1996) by POAM II (Fromm *et al.*, 1997). Figure 3-14 shows SAM II and POAM II zonally averaged PSC sighting probabilities for the Antarctic. Both datasets show Antarctic PSCs forming from mid-May until early November, with sighting probability increasing throughout the winter to a peak of about 60% in mid to late August. There is a downward trend in the altitude of peak sighting probability

related to the seasonal descent of the altitude of coldest temperature, a feature also noted in lidar measurements (David *et al.*, 1998). SAM II and POAM II Arctic PSC sighting probabilities (not shown) differ in an absolute sense, primarily because of the difference in latitude sampling of the two instruments. However, the Arctic data records are qualitatively similar, showing that PSCs occur there much less frequently and over a shorter time period (from early December to mid March) than in the Antarctic (Poole and Pitts, 1994).

Poole and Pitts (1994) noted that PSC existence temperatures inferred from SAM II data were nearly constant throughout the winter season in the Arctic, whereas a gradual lowering of existence temperatures was inferred over Antarctica from June through September, indicating a systematic removal of HNO₃ and H₂O by denitrification and dehydration in that region. The
existence temperatures in the early winter season in both hemispheres were close to the expected NAT equilibrium temperatures. On the other hand, gas-phase $HNO_3$ measurements by the Microwave Limb Sounder (MLS) instrument (Santee et al., 1998) show early-winter $HNO_3$ condensation beginning at roughly 3 K below $T_{NAT}$. This indicates that STS PSC particles were formed initially. Later in the season, $HNO_3$ concentrations approached the equilibrium levels for NAD (or NAT).

### 3.4.4.3 EFFECTS OF VOLCANIC LOADING ON PSC FORMATION

Large concentrations of volcanic SSA were observed to influence PSC formation in the years following the 1991 Mt. Pinatubo eruption. Antarctic measurements by Deshler et al. (1994b) showed that PSCs formed in the volcanic layer had numerous small particles ($r < 0.5$ µm) and essentially no particles with $r > 2$ µm, and that surface areas increased by a factor of 4 to 6 relative to the surrounding SSA. At altitudes above the volcanic layer, particles with $r \approx 2-3$ µm were observed, as had been the usual case in Antarctica prior to the eruption. Lidar observations show that most of the PSCs formed near the peak in SSA backscatter and at lower altitudes in the first years after eruption of Mt. Pinatubo. The larger volcanic aerosol loading apparently led to increased Type 1b PSC formation at higher temperatures, compared to pre-volcanic conditions (David et al., 1998). Satellite measurements also show a more pronounced $HNO_3$ uptake in STS droplets at higher temperatures under the enhanced aerosol loading from Mt. Pinatubo (Santee et al., 1998). Both of these findings are in good agreement with the equilibrium STS model calculations of Carslaw et al. (1994).

## 3.5 OUTSTANDING ISSUES IN STRATOSPHERIC PARTICLES

One outstanding issue that must be resolved is the quantification of the non-volcanic background SSA level, which cannot be determined from the current data record because of the influence of volcanoes. Understanding of this background level is required to assess the validity of models of SSA production, to allow detection of possible future anthropogenic trends, and to determine the partitioning of ozone destruction amongst the chlorine, hydrogen, and nitrogen radical families (Wennberg et al., 1994). Except for this issue, there is a general

consistency of SSA observations with dynamical and microphysical modeling results.

In contrast, the interpretation of PSC observations still suffers from many unknowns and uncertainties with relevance to stratospheric chemistry (see Chapter 7). Recent progress in understanding PSCs came with the recognition that some of these clouds are liquid and do not require any kind of phase transition during their growth. While there is overwhelming evidence for the existence of solid particles in the stratosphere, our understanding of the mechanisms of phase transitions leading to the formation of solid PSC particles is still very poor. Solid particle formation could be crucial for the heterogeneous chemical reactions occurring on PSCs because the solid $HNO_3$ hydrates can exist under warmer conditions than STS clouds. Furthermore, solidification of a few particles by some highly selective mechanism appears to be a necessary prerequisite for the denitrification of the polar vortices, which may lead to enhanced polar ozone destruction. For the modeling of Type 1a PSC formation, knowledge of the particle size distribution and phase at one particular point in time may not be sufficient, but rather the full thermal history of the air parcel containing the cloud particles may be required. There is a need for quasi-Lagrangian in situ measurements to follow particle evolution and to observe freezing events themselves.

Another outstanding issue is the effect that increases in source gases coupled with decreasing stratospheric temperatures may have on stratospheric particles. There are reports of an upward trend in $H_2O$, and Chapter 5 presents clear evidence of a long-term cooling in the lower stratosphere. An increase in $HNO_3$ in the lower stratosphere might also be expected because there is an upward trend in tropospheric $N_2O$. Both the surface area and heterogeneous reactivity of SSA should increase under these conditions. Surface area increases should be small, but there may be larger increases in heterogeneous reactivity (see Chapter 7).

More significant changes may occur in the properties and heterogeneous reactivity of Type 1b PSCs. For example, Figure 3-1 illustrates that particle surface area increases by 5-18% per 0.1 K temperature decrease in the interval between 192 and 191 K, around the STS threshold temperature. This could be of particular significance in the Arctic, where PSCs often form in limited synoptic-scale cold-temperature areas. The STS threshold temperature is also shifted upward by approximately 1 K per 1 ppm increase in $H_2O$ in the lower strato-
sphere. An example of this sensitivity is shown by HALOE data late in the 1993 Antarctic spring, where there was enhanced PSC development due to an intrusion of H₂O from lower latitudes (Hervig *et al.*, 1997). However, no trend in PSCs is discernible in the present satellite PSC data record because of the relatively short record length and the large interannual variability in the number of clouds observed.

It also should be noted that current models used for assessment calculations are limited in the accuracy to which they can represent PSC processes. For example, only large-scale PSC events can be reproduced in global models because of the models' relatively coarse resolution. Furthermore, most models include relatively simple microphysical assumptions or parameterizations, whereas field observations show that PSCs exhibit very complex and variable behavior. 2-D models inherently cannot represent the fact that PSC formation is not zonally symmetric, particularly in the Arctic, as shown by SAM II and POAM II data. Some progress has been made in this area (Considine *et al.*, 1994), whereby PSC formation is modeled using statistical temperature fluctuations about the zonal mean values, hence producing more PSCs than would occur using zonal average temperatures. Even a 3-D formulation has limitations in capturing small-scale phenomena such as gravity or mountain waves, where rapid temperature fluctuations may result in nonequilibrium microphysical processes that could be important in forming solid PSC particles, especially in the Arctic. The integral role of such phenomena in ozone depletion is unclear.

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