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184 Atmospheric Chemistry

days to a week. An important exception is COS,

which is very stable in the troposphere and, as a

result, has a relatively long residence time (2 years)

and a large and relatively uniform concentration

(0.5 ppbv, compared to 0.1 ppbv for DMS, 0.2

ppbv for H

S and SO

, and 0.05 ppbv for CS

Consequently, COS is the most abundant sulfur com-

pound in the troposphere. However, because it is

relatively unreactive, it is generally ignored in tropos-

pheric chemistry. The relatively long residence time

of COS enables it to be mixed into the stratosphere

gradually, where, converted by UV radiation, it is

the dominant source of the sulfate aerosol during

volcanically quiescent periods (see Section 5.7.3).

The flux of anthropogenic emissions of sulfur to

the atmosphere is known quite well; it is 78 Tg(S)

per year, which is greater than estimates of the

natural emissions of reduced sulfur gases to the

atmosphere (Fig. 5.15). Therefore, the global sulfur

budget is significantly affected by human activities.

Anthropogenic emissions of sulfur are almost entirely

in the form of SO

and 90% are from the northern

hemisphere. The main sources are the burning of coal

and the smelting of sulfide ores.

The main mechanisms for removing sulfur from the

atmosphere are wet and dry deposition. For example,

of the 80 Tg(S) per year of SO

that is oxidized to

about 70 Tg(S) per year occurs in clouds,

which is subsequently wet deposited; the remainder

2

is oxidized by gas-phase reactions and dry deposited

(Fig. 5.15).

Exercise 5.9 Using the information given in Fig. 5.15,

estimate the residence time of SO

in the troposphere

with respect to influxes.

Solution: From Fig. 5.15 we see that the magnitude

of the tropospheric reservoir of SO

is 0.3 Tg(S).

By adding the influxes of SO

shown in Fig. 5.15,

the total influx is found to be 120.5 Tg(S) per year.

Therefore,

residence time of SO

in the troposphere

■

5.7 Stratospheric Chemistry

In passing across the tropopause, from the tropo-

sphere to the stratosphere, there is generally an

abrupt change in concentrations of several important

trace constituents of the atmosphere. For example,

within the first few kilometers above the tropopause,

water vapor decreases and O

concentrations often



0.3

120.5

 2  10

3

years  1 day



magnitude of

tropospheric reservoir

total influx to troposphere

0.2

To the

stratosphere

0.3

0.9

0.5

0.3

0.5

0.7

4075103

CONTINENTS

Soils and

marshlands

10 73

(

2–

Volcanoes

Dry

deposition

Wet

deposition

Biomass

burning

Volcanoes

Anthropogenic

(fossil fuel

combustion)

Dry

deposition

Soils,

plants

and

industry

Biogenic Biogenic Biogenic

Soils Dry

deposition

[0.03 Tg(S)]

[1.2 Tg(S)]

[0.3 Tg(S)]

Sea-

salt

40–320

[0.1 Tg(S)]

DMS

[0.05 Tg(S)]

COS

[2.5 Tg(S)]

2–

CONTINENTS

[2 x 10

Tg(S)]

OCEANS

[1.6 x 10

Tg(S)]

Fig. 5.15 As for Fig. 5.14 but for sulfur-containing species in the troposphere. Fluxes are in Tg(S) per year. For clarity, wet and dry

removal are shown only over the continents, although they also occur over the oceans. [Adapted from P. V. Hobbs, Introduction to

Atmospheric Chemistry, Camb. Univ. Press, 2000, p. 150. Reprinted with the permission of Cambridge University Press.]

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5.7 Stratospheric Chemistry 185

increase by an order of magnitude. The strong verti-

cal gradients across the tropopause reflect the fact

that there is very little vertical mixing between the

relatively moist, ozone-poor troposphere and the dry,

ozone-rich stratosphere. Within the stratosphere the

air is generally neutral or stable with respect to verti-

cal motions. Also, the removal of aerosols and trace

gases by precipitation, which is a powerful cleansing

mechanism in the troposphere, is generally absent in

the stratosphere. Consequently, materials that enter

the stratosphere (e.g., volcanic effluents, anthro-

pogenic chemicals that diffuse across the tropopause

or are carried across the tropopause by strong

updrafts in deep thunderstorms, and effluents from

aircraft) can remain there for long periods of time,

often as stratified layers.

This section discusses three topics of particular

interest in stratospheric chemistry: unperturbed

(i.e., natural) stratospheric O

, anthropogenic pertur-

bations to stratospheric O

, and sulfur in the strato-

sphere. The emphasis in this chapter is on chemical

processes. However, it should be kept in mind that

chemical, physical, and dynamical processes in the

atmosphere are often intimately entwined.

5.7.1 Unperturbed Stratospheric Ozone

a. Distribution

Ozone in the stratosphere is of great importance for

the following reasons:

• It forms a protective shield that reduces the

intensity of UV radiation (with wavelengths

between 0.23 and 0.32



27,28

from the sun

that reaches the Earth’s surface.

• Because of the absorption of UV radiation

by O

, it determines the vertical profile of

temperature in the stratosphere.

• It is involved in many stratospheric chemical

reactions.

In 1881 Hartley

measured the UV radiation

reaching the Earth’s surface and found a sharp cutoff



 0.30



m; he correctly attributed this to absorp-

tion by O

in the stratosphere. Subsequent ground-

based UV measurements as a function of the sun’s

elevation and the first measurements of the concen-

trations of stratospheric O

by balloons in the 1930s

placed the maximum O

concentration in the lower

part of the stratosphere.

Figure 5.16 shows results of more recent measure-

ments of O

.The presence of an ozone layer between

heights of 15–30 km is clear. However, the O

layer

is highly variable: its height and intensity change

with latitude, season, and meteorological conditions.

Electromagnetic radiation in this wavelength range is dangerous to living cells. Radiation with a wavelength (



)  0.29



m is lethal

to lower organisms and to the cells of higher organisms. Radiation with



 0.290–0.320



m (UV-B radiation), which causes sunburn,

has adverse effects on human health and on animals and plants. Were it not for O

in the stratosphere, radiation from the sun with



 0.23–0.32



m, would reach the Earth’s surface unhindered. Ozone strongly absorbs UV radiation in just this wavelength band. For this

reason, less than 1 part in 10

of the flux of solar radiation at the top of the atmosphere with



 0.25



m reaches the Earth’s surface.

As shown in Chapter 2, the absorption of UV radiation by O

was essential for the emergence of life (as we know it) on Earth.

It is common to use nanometers (nm; 1



m  10

nm) as the unit of wavelength in the ultraviolet and visible regions. However, for

consistency with earlier chapters, we use micrometers (



m).

W. N. Hartley (1846–1913) Irish spectroscopist. Professor of chemistry at Royal College of Science, Dublin.

Altitude (km)

Ozone concentration (molecules cm

–3

)

250 325

400

450

30°

43°

59°

DU DU DU

9°

Fig. 5.16 Mean vertical distributions of ozone concentra-

tions based on measurements at different latitudes (given

in degrees). Note the increase in the total ozone column

abundance (given in DU) with increasing latitude. [Adapted

from G. Brasseur and S. Solomon, Aeronomy of the Middle

Atmosphere, D. Reidel Pub. Co., 1984, Fig. 5.7, p. 215.

Springer Science and Business Media.]

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186 Atmospheric Chemistry

Figure 5.16 includes the total ozone column abun-

dance in Dobson units (DU)—see Footnote 13 in this

chapter for the definition of DU.

The greatest column densities of O

in the northern

hemisphere occur in polar latitudes in spring; in the

southern hemisphere the spring maximum is at mid-

latitudes. Because O

is produced by photochemical

reactions, the production is a maximum in the strato-

sphere over the tropics. The peaks in concentrations

at polar and midlatitudes are attributed to meridional

and downward transport of O

away from the equa-

tor, although at any given point in the atmosphere

the balance between the production and loss of O

and its flux divergence, determines the O

concentra-

tion at that point. It is clear from Fig. 5.16 that much

of the meridional contrast in the total column abun-

dance of O

is due to differences in the profiles below

20 km, which are largely determined by transport.

Since 1960 remote sensing measurements from satel-

lites (see Box 5.5) have provided a wealth of informa-

tion on the global distribution of O

and the variations

in its vertical profiles and column abundance.

In situ measurements of vertical profiles of O

may

be obtained from ozonesondes, which can be car-

ried on radiosonde balloons. The O

sensor consists

of two electrolytic cells, each containing a solution

of potassium iodide (KI). The cells are initially in

chemical and electrical equilibrium. However,

when an air sample containing O

is drawn through

one of the cells, the equilibrium is perturbed and an

electric current flows between the cells. The amount

of electric charge, which is proportional to the par-

tial pressure of the O

in the ambient air, is contin-

uously transmitted to a ground station along with

the ambient pressure and temperature. In this way,

a vertical profile of O

is obtained, the integration

of which provides the O

column from ground level

up to the height of the balloon.

The O

column from the ground to the top of

the atmosphere can be measured by passive remote

sensing using a Dobson spectrophotometer.This

is done by measuring the amount of UV sunlight

that reaches the ground and deducing from

this how much UV absorption occurred due to

. Absorption by O

occurs in the UV-B region

(



 0.290–0.320



m). However, clouds and some

aerosol particles also absorb in this wavelength

band. Therefore, a region of the electromagnetic

spectrum where O

absorbs only weakly, but

clouds and aerosol absorb similarly to the UV-B

region, is also monitored. By ratioing the two

measured values, absorption by O

in the total

vertical column can be obtained.

Ozone can be derived from satellite observations

using any of the four passive techniques depicted

in Fig. 5.17, namely by backscatter UV(BUV),

occultation, limb emission, and limb scattering.

For determination of the total O

column by the

BUV technique (Fig. 5.17a), two pairs of measure-

ments are needed: the incoming UV irradiance and

the backscattered UV radiance at a wavelength that

is strongly absorbed by O

and a similar pair of

measurements at a wavelength that is absorbed only

weakly by O

.The difference between these two

pairs of measurements can be used to infer the total

column. Vertical profiles of O

can be obtained

using the BUV technique. Because O

absorbs more

strongly at shorter wavelengths, solar radiation with

progressively shorter wavelengths is absorbed at

progressively higher altitudes. Therefore, the radia-

tion at a particular wavelength in the UV can be

scattered only above a certain height. By measuring

backscattered radiation at a number of wavelengths,

the vertical profile of O

can be deduced.

Occultation instruments measure radiation

through the limb of the atmosphere (see Fig. 5.17b)

when the sun, moon, or a star is rising or setting.

From measurements of the amounts of radiation at

various wavelengths absorbed by the atmosphere,

vertical profiles of various trace constituents

can be derived. This technique has been used to

monitor O

since 1984.

In the limb emission technique, the concentration

of O

is derived from measurements of infrared or

microwave radiation emitted by the atmosphere

along the line of sight of the instrument (Fig. 5.17c).

Limb infrared emissions can be used to derive tem-

perature, O

O(g), CH

O, NO

, CFCs, and

the location of polar stratospheric clouds in the

upper atmosphere. The advantage of utilizing

microwave emission is that it passes through clouds

and therefore can provide measurements lower in

5.5 Techniques for Measuring Atmospheric Ozone

Continued on next page

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5.7 Stratospheric Chemistry 187

Fig. 5.17 Four passive techniques for measuring ozone from satellites. [Courtesy of P. Newman, NASA Goddard Space

Flight Center.]

the atmosphere than infrared emissions. However,

microwave detectors are large, heavy, and require

more power than infrared detectors.

Limb scattering (Fig. 5.17d) employs aspects

of the three passive remote sensing techniques

discussed briefly earlier. However, it measures

scattered light rather than light directly from the

sun or moon. Consequently, if the sun is visible,

this technique can provide essentially continuous

measurements. Limb scattering works best with

, but H

O(g), NO

,SO

, and aerosols can also

be measured.

Ozone profiles (as well as those of other trace

constituents) can be obtained using differential

absorption lidar (DIAL), which is an active remote

sensing technique. The radiation source is a power-

ful lidar. A telescope, generally located near the

laser, is used to collect the radiation backscattered

by the atmosphere along the line of sight of the lidar

beam. The time interval between transmission of a

pulse and its return can be used to determine the

range of the scatterers. The attenuation of the beam

provides information on atmospheric trace con-

stituents. Differences between transmitted and

returned wavelengths produced by the Doppler

5.5 Continued

effect can be used to infer atmospheric motions. As

in the BUV technique, the ratio of two returned sig-

nals can be used to deduce the concentration of the

species. This technique can provide measurements

with high spatial resolution in the vertical. Because

DIAL instruments are large and heavy, they have

generally been used only from the ground.

The advantage of satellite-born instruments

is their extensive spatial coverage over relatively

long time periods. Disadvantages are that the

measurements of trace constituents are not direct

(they have to be derived from radiation measure-

ments) and typically they do not have very high

spatial resolution. Remote sensing measurements

from the ground can be continuous, but each

instrument is generally located at one site.

Remote sensing measurements from aircraft pro-

vide mobility and better spatial resolution in the

vertical. In situ measurements (from aircraft and

balloons) are generally more direct, they can be

related to chemical standards, and they can pro-

vide high spatial resolution. However, they are

inherently very intermittent. Airborne in situ

measurements have been used extensively to vali-

date satellite measurements.

(a) Backscatter ultraviolet (b) Occultation

satellite

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188 Atmospheric Chemistry

b. Chapman’s theory

In 1930 Chapman

proposed a simple chemical

scheme for maintaining steady-state concentrations

of O

in an “oxygen-only” (i.e., O, O

, and O

) strato-

sphere. The reactions were the dissociation of O

solar UV radiation (at



 0.242



(5.43)

the reaction of atomic oxygen and molecular oxygen

to form O

(5.44)

(where M represents N

and O

), the photodissociation

of O

(which occurs for



 0.366



(5.45)

and the combination of atomic oxygen and O

form O

(5.46)

Reactions (5.43)–(5.46) are called the Chapman

reactions.The rate coefficients (j’s and k’s) are tem-

perature dependent.

The O atoms produced by (5.43) undergo numerous

collisions with N

and O

molecules, most of which do

not result in any permanent products. Occasionally,

however, a three-body collision occurs [as represented

by the left side of (5.44)] in which M absorbs the excess

energy of the collision and a stable O

molecule is pro-

duced. Note that the excess energy acquired by M,

which is in the form of thermal energy that warms the

stratosphere, derives primarily from the energy of

the photon in (5.43), i.e., from solar energy. Because

the concentration of M decreases with increasing

altitude, the time constant for converting O atoms and

molecules to O

by (5.44) increases with altitude

(e.g., from a few seconds at 40 km to 100 s at 50 km).

After it is formed, the O

molecule remains intact

until it is split apart by the photolytic reaction (5.45).

Reactions (5.44) and (5.45) are fast and they continu-

ally cycle oxygen atoms back and forth between O and

. Ozone molecules are removed from the system by

(5.46). The time constant for this reaction is 1 day

at 40 km and a fraction of a day at 50 km (where the

O  O

 h



 O

O  O

 M :

 M

 h



concentration of O atoms is greater). This reaction

can take place as given by (5.46) or it can be catalyt-

ically accelerated (e.g., by chlorine or bromine), as

discussed in Section 5.7.1c. A simplified schematic of

the Chapman reactions is shown in Fig. 5.18.

Stratospheric O

concentrations exhibit minor diur-

nal variations with time of day. After sunset, both the

source and the sink of O

(5.43) and (5.45), respect-

ively, are switched off, and the remaining O atoms

are then converted to O

within a minute or so by

(5.44). When the sun rises, some of the O

molecules

are destroyed by (5.45) but they are reformed by

(5.43) followed by (5.44).

Exercise 5.10 During the daytime in the strato-

sphere a steady-state concentration of atomic oxygen

may be assumed. Use this fact to derive an expression

for dn

dt in terms of n

, n

, k

, and j

, where n

,and n

are the concentrations of atomic oxygen,

molecular oxygen, and ozone, respectively. Assuming

that the removal of atomic oxygen by (5.46) is small,

and (at 30 km) j

 j

, derive an expression

for the steady-state concentration of atomic oxygen

in the stratosphere at 30 km during the daytime in

terms of n

, n

, k

,and j

,where n

is the concen-

tration of M in (5.44). Hence, derive an expression

for dn

dt in terms of n

, n

, j

, k

, and k

Solution: From (5.43)–(5.46)

(5.47)

If atomic oxygen is in steady-state dn

dt  0, and

(5.48)2j

 j

 k

 k

 2j

 k

 j

 k

O + O

hν

+ M

ODD

OXYGEN

fast

slow

hν

Fig. 5.18 Schematic of the Chapman reactions. [Adapted

from The Chemistry and Physics of Stratospheric Ozone, A. Dessler,

Academic Press, p. 44, copyright (2000), with permission

from Elsevier.]

Sydney Chapman (1888–1970) English geophysicist. Chapman made important contributions to a wide range of geophysical

problems, including geomagnetism, space physics, photochemistry, and diffusion and convection in the atmosphere.

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5.7 Stratospheric Chemistry 189

From (5.44)–(5.46)

(5.49)

From (5.48) and (5.49)

(5.50)

If the removal of atomic oxygen by (5.46) is small, we

have from (5.47)

(5.51)

Therefore, since dn

dt  0 and j

 2j

(5.52)

Substituting (5.52) into (5.50) yields

(5.53)

■

The Chapman reactions reproduce some of the

broad features of the vertical distribution of O

in the stratosphere. For example, they predict that

concentrations should reach maximum values at

an altitude of 25 km. Indeed, prior to 1964 it was

generally believed that the Chapman theory pro-

vided an adequate description of stratospheric ozone

chemistry. This is now known not to be the case.

For example, although the Chapman reactions pre-

dict the correct shape for the vertical profile of O

they overpredict the concentrations of stratospheric

by about a factor of two in the tropics and

they underpredict O

concentration in middle to

high latitudes. Also, model calculations based on the

Chapman reactions predict that the global rate of pro-

duction of O

in spring is 4.86  10

molecules s

1

(of which 0.06  10

molecules s

1

are transported

to the troposphere). However, the loss of “odd oxy-

gen” is only 0.89  10

molecules s

1

.This leaves a

net 3.91  10

molecules s

1

, which would double

atmospheric O

concentrations in just 2 weeks.

Because O

concentrations are not increasing, there

must be important sinks of odd oxygen in the strato-

 2j





 2j

 k

 j

 2j

 2k

 k

 j

 k

sphere in addition to reaction (5.46). As shown later,

there are catalytic cycles involving nitrogen com-

pounds, H, OH, Cl, and Br that deplete O

in the

stratosphere. Also, there is an equator-to-pole circu-

lation in the stratosphere, sometimes referred to as

the Brewer–Dobson circulation, which transports O

from its primary source in the tropical stratosphere

poleward and downward into the extratropical lower

stratosphere.

c. Catalytic chemical cycles

Most of the catalytic reactions that have been pro-

posed for the removal of stratospheric odd oxygen

are of the form

(5.54a)

(5.54b)

Net: (5.55)

where X represents the catalyst and XO the inter-

mediate product. Provided that reactions (5.54) are

fast, reaction (5.55) can proceed much faster than

reaction (5.46). Also, because X is consumed in

(5.54a) but regenerated in reaction (5.54b) and pro-

vided there is no appreciable sink for X, just a few

molecules of X have the potential to eliminate large

numbers of O

molecules and atomic oxygen.

The first candidate that was suggested for species

X in (5.54) in the upper atmosphere (actually in the

mesosphere where excited atomic oxygen is common)

was OH, which has at least three sources

(5.56)

(5.57)

and

(5.58)

The H

O and CH

in (5.56) and (5.57) ascend into

the stratosphere in the tropics. During the passage

through the cold tropical tropopause, most of the

water vapor condenses out, which accounts for the

 O* : H  OH

 O* : CH

 OH

O  O* : 2OH

O  O

: 2O

XO  O : X  O

X  O

: XO  O

In the Chapman reactions, odd oxygen is produced only by reaction (5.43) and is lost only in reaction (5.46). Reactions (5.44) and

(5.45) intraconvert atomic oxygen and O

and determine the ratio [O][O

]; both reactions are fast during the day. Therefore, atomic oxygen

and O

are interconverted rapidly, as shown in Fig. 5.18. Below about 45 km in the stratosphere, O

accounts for 99% of the odd oxygen.

P732951-Ch05.qxd 12/09/2005 09:05 PM Page 189

190 Atmospheric Chemistry

dryness of the stratosphere. However, CH

is not

affected by the low temperatures and is therefore

lofted into the lower stratosphere in concentrations

similar to those in the upper troposphere.

Below 40 km the OH from (5.56)–(5.58) acts as

catalyst X in (5.54) to destroy odd oxygen

(5.59a)

(5.59b)

Net: (5.60)

Below 30 km, where atomic oxygen is very scarce,

the following catalytic cycle, which destroys O

,is

increasingly important

(5.61a)

(5.61b)

Net: (5.62)

A catalytic cycle that is important for the destruc-

tion of O

in the middle stratosphere is

(5.63a)

(5.63b)

Net: (5.64)

where NO has replaced X in (5.54). Nitric oxide

(NO) is produced in the stratosphere by

(5.65)

At a temperature of 53 °C (which is typical of

the stratosphere), the rate coefficients for reactions

(5.63a) and (5.63b) are 3.5  10

21

and 9.3 

18

molecules

1

, respectively, compared

to 6.8  10

22

molecules

1

for k

in (5.46).

Whether reactions (5.63a) and (5.63b) destroy odd

oxygen faster than reaction (5.46) depends on the

concentrations of NO

and O

In the 1970s it was suggested that catalytic cycles

involving reactive chlorine (Cl) and bromine (Br)

could be important in the destruction of O

.The

natural sources of Cl and Br in the stratosphere

are the destruction of methyl chloride (CH

Cl) and

methyl bromide (CH

Br), respectively. Chlorine and

Br can then serve as X in (5.54), with intermediates

O  O* : 2NO

O  O

: 2O

 O : NO  O

NO  O

: NO

 O

: 3O

 O

: OH  2O

OH  O

: HO

 O

O  O

: 2O

 O : OH  O

OH  O

: HO

 O

ClO and BrO, respectively. The following cycle of

reactions, which couple these two intermediates,

(5.66a)

(5.66b)

(5.66c)

Net: (5.67)

is a more powerful mechanism for destroying O

the lower stratosphere than those discussed previ-

ously because O

is much more abundant than atomic

oxygen. Because the stable reservoirs for Br and BrO

(i.e., BrONO

and HOBr) are not as stable as the

reservoirs for Cl and ClO (i.e., HCl and ClONO

) on

a per molecule basis, Br is more efficient in destroy-

ing O

than Cl.

The destruction of stratospheric O

by the various

chemical mechanisms discussed earlier, and many

other possible mechanisms not discussed here are

not simply additive because, as illustrated by (5.66),

the species in one cycle can react with those in

another cycle. Therefore, advanced numerical models

that consider all of the known suspects and that

utilize measurements of species concentrations and

temperature-dependent rate coefficients must be used

to unravel the relative importance of the various

mechanisms for destroying O

in the stratosphere.

This task is complicated further by the fact that

concentrations of many of the prime suspects are

changing with time due to anthropogenic emissions

(see Section 5.7.2).

5.7.2 Anthropogenic Perturbations to

Stratospheric Ozone: The Ozone Hole

If the concentration of catalyst X in (5.54) is

increased significantly by human activities, the bal-

ance between the sources and sinks of atmospheric

will be disturbed and stratospheric O

concen-

trations can be expected to decrease. One of the

first concerns in this respect was raised in response

to a proposal in the 1970s to build a fleet of super-

sonic aircraft flying in the stratosphere. It was noted

that aircraft engines emit nitric oxide (NO), which

can decrease odd oxygen by the cycle (5.63). The

proposal was eventually rejected on both envi-

ronmental and economic grounds. At the time of

writing (2005), there are not sufficient aircraft flying

: 2O

Cl  O

: ClO  O

Br  O

: BrO  O

BrO  ClO : Br  Cl  O

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5.7 Stratospheric Chemistry 191

in the stratosphere to significantly perturb stratos-

pheric O

Of much greater concern, with documented

impacts, is the catalytic action of chlorine, from indus-

trially manufactured chlorofluorocarbons (CFCs),

in depleting stratospheric ozone.

CFCs are com-

pounds containing Cl, F, and C; CFC-11 (CFCl

)

and CFC-12 (CF

) are the most common.

CFCs

were first synthesized in 1928 as the result of a search

for a nontoxic, nonflammable refrigerant. Over the

next half-century they were marketed under the trade

name Freon and became widely used, not only as

refrigerants, but as propellants in aerosol cans, inflat-

ing agents in foam materials, solvents, and cleansing

agents. Concern about their effects on the atmos-

phere began in 1973 when it was found that CFCs

were spreading globally and, because of their inert-

ness, were expected to have residence times ranging

up to several hundred years in the troposphere.

Such long-lived compounds eventually find their way

into the stratosphere

where, at altitudes 20 km,

they absorb UV radiation in the wavelength interval

0.19–0.22



m and photodissociate

(5.68)

and

(5.69)

The chlorine atom released by these reactions can

serve as the catalyst X in cycle (5.54), thereby

destroying odd oxygen in the cycle

(5.70a)

(5.70b)

Net: (5.71) O

 O : 2O

ClO  O : Cl  O

Cl  O

: ClO  O

 h



: CF

Cl  Cl

CFCl

 h



: CFCl

 Cl

In 1990, 85% of the chlorine in the stratosphere

originated from anthropogenic sources. Because

CFCs absorb strongly in the infrared, they are also

significant greenhouse gases.

The first evidence of a significant depletion in

stratospheric O

produced by anthropogenic chemicals

in the stratosphere came, surprisingly, from measure-

ments in the Antarctic. In 1985, British scientists, who

had been making ground-based, remote sensing meas-

urements of O

at Halley Bay (76 °S) in the Antarctic

for many years, reported a 30% decrease in springtime

(october) total column O

since 1977. These observa-

tions were subsequently confirmed by remote sensing

measurements from satellite and by airborne in situ

measurements. Shown in Fig. 5.19 are a series of satel-

lite measurements of the total O

column in the south-

ern hemisphere. The high O

values (red and orange

colors) that encircle the Antarctic continent and the

lower O

values (green) over the continent itself in the

1970s (upper four images) are considered to be close

to natural. The Antarctic ozone hole, as it came to be

called, is apparent in the lower four images in Fig. 5.19

as the anomalously low O

values (purple) over the

continent in October from 1999 to 2002.

The areal extent of the Antarctic ozone hole,

derived from satellite measurements for the period

September 7 to October 13 1979–2003, are shown in

Fig. 5.20. From 1979 to 2001 the ozone hole grew pro-

gressively until it occupied an area (25 million km

)

similar to that of North America. However, in 2002

the area occupied by the ozone hole decreased dra-

matically, but returned to its former size in 2003; the

reasons for these dramatic changes are discussed

later in this chapter.

The presence of an ozone hole over the Antarctic

raised several intriguing questions. Why over the

Antarctic? Why during spring? Why was the magnitude

of the measured decreases in O

much greater than

predictions based solely on gas-phase chemistry of the

In 1995, Paul Crutzen, Mario Molina, and Sherwood Rowland were awarded the Nobel Prize in chemistry for predicting stratos-

pheric O

loss by CFCs and nitrogen-containing gases.

Compounds containing H, F, and C are called hydrofluorocarbons (HFCs), and compounds containing H, Cl, F, and C are called

hydrochlorofluorocarbons (HCFCs). HFCs and HCFCs are considered to be more “environmentally friendly” than CFCs because they are

partially destroyed by OH in the troposphere.

Long-lived compounds enter the stratosphere primarily in air that ascends through the tropical tropopause. After entering

the lower stratosphere the compounds are carried by the Brewer–Dobson circulation to higher altitudes (where they are broken

down photolytically) or to higher latitudes. On average, it takes a year or two for a CFC molecule to move from the tropical tropo-

sphere to the upper stratosphere. However, it takes several decades for all of the air in the troposphere to be cycled through the upper

stratosphere.

P732951-Ch05.qxd 12/09/2005 09:05 PM Page 191

192 Atmospheric Chemistry

type outlined earlier? Why was the size of the ozone

hole so much less in 2002 than in other recent years?

The research that led to the answers to these questions,

which is outlined later, was one of the triumphs of

20th century science.

During the austral winter (June–September),

stratospheric air over the Antarctic continent is

restricted from interacting with air from lower lati-

tudes by a large-scale vortex circulation, which is

bounded at its perimeter (called the vortex collar) by

strong westerly winds encircling the pole. Because of

the lack of solar heating in the austral winter, the air

within the vortex is extremely cold and it lies within

the sinking branch of the Brewer–Dobson circulation

(Fig. 5.21). High-level clouds, called polar stratospheric

clouds (PSCs), form in the cold core of the vortex,

where temperatures can fall below 80 °C.

In the

austral spring, as temperatures rise, the winds around

the vortex weaken, and by late December the vortex

disappears. However, during winter, the vortex serves

as a giant and relatively isolated chemical reactor

in which unique chemistry can occur. For example,

Oct. 00 Oct. 01 Oct. 02Oct. 99

Oct. 71 Oct. 72 Oct. 79Oct. 70

Total Ozone (DU)

500.

120.

Fig. 5.19 Satellite observations of the total ozone column in the southern hemisphere during October for 8 years from 1970 to

2002. The color scale is in Dobson units (DU). The small black circles over the pole in 1970, 1971, and 1972 are a reflection of

missing data. [Courtesy P. Newman, NASA Goddard Space Flight Center.]

1980 1985 1990 1995 2000

Size (Million km

)

Antarctica area

North America area

Fig. 5.20 Average areal extent of the ozone hole (defined

as less than 220 Dobson units) averaged from September

7–October 13 of each year from 1979 to 2003. Vertical red

bars are the range of values for each period. Year-to-year

variations are caused by temperature fluctuations near the

polar vortex edge. Warmer years have smaller holes (e.g.,

2002), whereas colder years have larger ozone holes

(e.g., 2003). The blue horizontal line highlights the total area

of the Antarctic continent; for comparison, the area of North

America (dashed blue line) is about twice the area of

Antarctica. (Courtesy of P. Newman, NASA Goddard Space

Flight Center.)

During the drift of the Deutschland in the Weddell Sea in 1912, PSCs were observed in the late winter and early spring. Also, the

Norwegian–British–Swedish 1949–1952 expedition to the Antarctic frequently observed a thin “cloud-veil” in the lower stratosphere in

winter months. However, the widespread nature and periodicity of PSC formation in the Antarctic were not fully appreciated until they

were observed by satellites starting in 1979.

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5.7 Stratospheric Chemistry 193

although the concentrations of O

in the vortex

are normal in August, the concentrations of ClO in the

vortex are 10 times greater than just outside the vor-

tex collar. In September, when sunlight returns to the

polar cap region, O

concentrations within the vortex

decrease dramatically. There are also sharp decreases

in the concentrations of oxides of nitrogen and water

vapor when passing from the outside to the inside of

the vortex collar. These decreases are due, respectively,

to the formation of nitric acid (HNO

) and to the con-

densation of water at the very low temperatures inside

the vortex. These two condensates form three types of

PSCs. Type I PSCs, which condense near 78 °C, prob-

ably consist of a mixture of liquid and solid particles of

nitric acid trihydrate (HNO

—NAT for short),

water and sulfuric acid. These particles are 1



in diameter, so they settle out very slowly (10 m

per day). Type II PSCs, which form near 85 °C, con-

sist of a mixture of ice and water together with some

dissolved HNO

. Because these particles are 10



in diameter, they settle out with appreciable speeds

(1.5 km per day). Type III PSCs (“mother-of-pearl”

clouds) are produced by the rapid freezing of con-

densed water in air flow over topography. However,

they are of limited extent and duration and do not

form over the South Pole.

As the particles in PSCs slowly sink, they remove

both water and nitrogen compounds from the strato-

sphere, processes referred to as dehydration and

denitrification, respectively. As we will see, the chemi-

cal reactions involved in the removal of these two

species play an important role in depleting O

in the

Antarctic vortex.

Although (5.70) is important in destroying O

the middle and upper stratosphere, it is less important

in the lower stratosphere (accounting for 5–25%

of the total loss of halogens there) because the con-

centration of atomic oxygen decreases with decreas-

ing altitude in the stratosphere and because most of

the Cl and ClO released into the stratosphere by

(5.68)–(5.70) are quickly tied up in reservoirs as HCl

and chlorine nitrate ClONO

(5.72)

and

(5.73)

Liberation of the active Cl atoms from the reservoir

species HCl and ClONO

is generally slow. However,

on the surface of the ice particles that form PSCs, the

ClO  NO

 M : ClONO

 M

Cl  CH

: HCl  CH

HNO

ClONO

HOCl

, HOCl

HCl

During

polar night

Following

sunrise in spring

Antarctica

South

America

Tasmania

(

–

)

Formation of PSCs

Release of Cl

and HOCl

Formation

of ClO and Cl

(eqns. (5.84)–(5.87))

Catalytic depletion

of O

(eqns. (5.88)–(5.89))

~25 km ~30 hPa

(eqns. (5.79)–(5.83))

Denitrification and

dehydration

and (ClO)

(eqn. (5.88))

Fig. 5.21 Schematic of the polar vortex (blue) over Antarctica. Large arrows indicate cold descending air. The sequence of events

(1 through 4) leading to the Antarctic ozone hole is summarized. For clarity, bromine reactions are not shown. See text for details.

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