Drake G.W.F. (editor) Handbook of Atomic, Molecular, and Optical Physics

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1293

Applications o

85. Applications of Atomic and Molecular Physics

to Global Change

While there has been a general understanding

and appreciation of the science involved in both

global warming and stratospheric ozone depletion

by atmospheric scientists for some time, detailed

understanding and rigorous proof has often been

lacking. Over the last ten years, there have been

many advances made in ﬁlling in the details

and there will continue to be rapid advances in

the future. This means that any article or book

discussing this topic becomes out of date as soon as

it is written. Nevertheless several recent references

on these topics are recommended [85.1–3].

Atomic and molecular structure and spec-

troscopy, as well as collision processes involving

atoms, molecules, ions and electrons, are impor-

tant to the study of all planetary atmospheres. For

additional information on this topic, see Chapt. 84

in this volume.

85.1 Overview.............................................1293

85.1.1 Global Change Issues .................1293

85.1.2 Structure of the Earth’s

Atmosphere ............................. 1293

85.2 Atmospheric Models

and Data Needs ...................................1294

85.2.1 Modeling the Thermosphere

and Ionosphere ........................1294

85.2.2 Heating and Cooling Processes .... 1295

85.2.3 Atomic and Molecular Data Needs1295

85.3 Tropospheric Warming/

Upper Atmosphere Cooling ...................1295

85.3.1 Incoming and Outgoing

Energy Fluxes ...........................1295

85.3.2 Tropospheric “Global” Warming .. 1296

85.3.3 Upper Atmosphere Cooling ......... 1297

85.4 Stratospheric Ozone .............................1298

85.4.1 Production and Destruction ........1298

85.4.2 The Antarctic Ozone Hole ............ 1299

85.4.3 Arctic Ozone Loss .......................1300

85.4.4 Global Ozone Depletion..............1300

85.5 Atmospheric Measurements ..................1300

References .................................................. 1301

85.1 Overview

85.1.1 Global Change Issues

Over the last several decades there has been increasing

concern about the global environment and the effect of

human perturbations on it. This whole area, which in-

volves a wide range of scientiﬁc disciplines, has become

known as Global Change. Knowledge of processes tak-

ing place in the atmosphere, oceans, land masses, and

plant and animal populations, as well as the interactions

between these various earth-system components is es-

sential to an overall understanding of global change –

both natural and human-induced.

The processes of atomic and molecular physics ﬁnd

greatest application in the area of atmospheric global

change. The two major issues which have received sig-

niﬁcant attention in both the media and the scientiﬁc

literature are: (1) global warming, due to the buildup

of infrared-active gases; and (2) stratospheric ozone de-

pletion due to an enhancement of destructive catalytic

cycles. Although both of these problems are thought to

be caused by atmospheric pollutants due to industrial-

ized human society, the general problem of air pollution

and its direct effects on plant and animal populations

will not be addressed here.

85.1.2 Structure of the Earth’s Atmosphere

The vertical temperature structure of the earth’s at-

mosphereshowninFig.85.1 provides an important

nomenclature that is widely used [85.4]. The atmosphere

is divided into regions called “-spheres”, in which the

sign of the temperature gradient with respect to altitude,

dT/dz, is constant. The regions in which the temper-

ature gradient changes sign are called “-pauses”. The

precise altitude pertaining to each of these regions can

vary depending upon latitude and the time of year.

Part G 85

1294 Part G Applications

140

120

100

Altitude (km)

Temperature (K)

100 150 200 250 300 350 400

Ionosphere

Thermosphere

Mesopause

Stratopause

Mesosphere

Stratosphere

Troposphere

Equatorial tropopause

Polar tropopause

Fig. 85.1 Vertical temperature proﬁle of the atmosphere

In the troposphere, covering the range from 0 to

≈ 15 km above the earth, the temperature steadily de-

creases with altitude. This is the most complex region

of the atmosphere, as it interacts directly with plant and

animal life, land masses and the oceans. It is the region

in which weather occurs.

The change in sign of the temperature gradient at

the tropopause to a positive dT/dz in the stratosphere

is due to heating by absorption of solar ultraviolet

radiation which photodissociates O

and O

.Thestrato-

sphere, extending from ≈ 15 to 50 km above the earth,

contains the ozone layer which shields the earth’s sur-

face from harmful ultraviolet radiation in the range

of 280–320 nm.

At the stratopause, the heating processes have

become too weak to compete with the cooling pro-

cesses, and throughout the mesosphere, approximately

50–85 km, the temperature again decreases with in-

creasing altitude. Cooling processes, which will be

discussed in Sect. 85.2.2, involve collisional excitation

of molecular vibrational modes which decay by radiat-

ing to space. The coldest temperatures in the atmosphere

are found at the mesopause, where the temperature

gradient once again becomes positive. From approxi-

mately 70 km upward, a very diffuse plasma called the

ionosphere exists due to photoionization of atoms and

molecules by short wavelength (UV and EUV)solar

radiation.

Throughout the thermosphere, which extends from

approximately 90 km upward, heating occurs because

the atmosphere has become so thin that there are very

few collisions and thus inefﬁcient equilibration of the

highly translationally excited atoms and ions with the

molecular species which can radiate in the infrared.

This “bottleneck” for energy loss causes increased heat-

ing. In the thermosphere and ionosphere, the thermal

inertia is very small and there are huge temperature

variations, both diurnally, and with respect to solar

activity.

The densities are low enough in the thermosphere,

ionosphere, and mesosphere, that the primary processes

determining the chemical and physical characteristics

of these regions are two-body processes and “half-

collision events” discussed elsewhere in this volume:

dissociative recombination, photoionization, photodis-

sociation, charge transfer, and collisional excitation

of molecular rotation and vibration. As the altitude

decreases, the density increases. Then three-body in-

teractions, interactions on surfaces (of aerosols and

ices), and complex chemical cycles together with dy-

namical effects such as winds determine the chemical

and physical characteristics of the stratosphere and

troposphere.

85.2 Atmospheric Models and Data Needs

While models are absolutely essential to the study of

any system as complex as the earth’s atmosphere, they

play a particularly fundamental role in exploring global

change issues. Models not only provide predictions of

future changes, but also allow exploration of sensitivi-

ties to particular parameters. Comparing the results of

a model with observations ultimately tests and chal-

lenges scientiﬁc understanding. Of critical importance

is the atomic and molecular data which goes into the

models.

Generally, atmospheric models become increas-

ingly complex as altitude decreases. General Circulation

Models (GCMs), incorporating thousands of chemical

reactions, global wind patterns, and abundances of large

numbers of trace species, require supercomputers in

order to model aspects of the troposphere and strato-

sphere. Tropospheric chemistry and transport models,

such as GEOS-CHEM [85.5], model the sources, evo-

lution, transport, and sinks of pollution, as well as the

oxidative capacity of the troposphere.

Part G 85.2

Applications of Atomic and Molecular Physics to Global Change 85.3 Tropospheric Warming/Upper Atmosphere Cooling 1295

85.2.1 Modeling the Thermosphere

and Ionosphere

The data necessary for modeling of the thermosphere

and ionosphere are described here. The primary compo-

nents of such a model include: a solar spectrum of photon

ﬂuxes as a function of wavelength, concentrations

of neutral species, photoabsorption, photodissociation

and photoionization cross sections as a function of

wavelength. The computer code brings all these ele-

ments together, calculating opacity as the solar radiation

propagates downward through the atmosphere, keep-

ing track of ion production and electron production.

Additional steps are required to calculate the abun-

dances of trace species such as NO

, necessitating

the inclusion of all relevant ion-neutral reactions. Elec-

tron energy degradation can be tracked by including

inelastic collisions of electrons with ions, atoms, and

molecules.

The primary neutral species are N

,O,andHe.

Below about 100 km, the atmosphere is fully mixed by

turbulence in the ratio 78% N

,21%O

and 1% trace

species such as O

and CO

. Above 100 km, turbulence

dies out and the atmospheric species are in diffusive

equilibrium, distributed by their molecular weight, with

atomic oxygen dominating above ≈ 150 km and He

dominating much higher. The major ions are N

and NO

85.2.2 Heating and Cooling Processes

In the upper atmosphere, heating occurs through ab-

sorption of short wavelength solar radiation to produce

ionization and dissociation, and is mediated by colli-

sions between electrons, ions, and neutrals. Ions and

electrons are created during the daytime and to a great

extent disappear during the night with the absence of

solar radiation. Processes such as dissociative recombi-

nation, the primary electron loss mechanism, heat the

gas:

e + O

→ O(

P) + O(

P) + 7eV, (85.1)

e + N

→ N(

S) + N(

S) + 6eV. (85.2)

Cooling takes place when the kinetic energy of the

gas is transformed through collisions into internal en-

ergy which can then be radiated away. The primary

coolant above ≈ 200 km is the ﬁne structure transition

of atomic oxygen, O(

)

hν

→ O(

), which is excited

by thermal collisions and radiates at 63 µm. From ap-

proximately 120 km to 200 km, the fundamental band of

NO, v = 1 → v = 0, which is excited by collisions with

atomic oxygen and radiates at 5.3 µm, dominates the

cooling. Below 120 km and throughout the mesosphere

and stratosphere, the primary coolant is the ν

band of

radiating at 15 µm. This transition is excited by col-

lisions of CO

and atomic oxygen. Cooling throughout

most of the atmosphere is accomplished through trace

species because the major molecular species, N

and O

are not infrared active.

85.2.3 Atomic and Molecular Data Needs

Knowledge of rate coefﬁcients for ion-neutral and

neutral-neutral reactions as a function of vibrational and

rotational excitation of the reactants is becoming in-

creasingly important, as there is recent evidence of more

internal excitation of molecular species than had previ-

ously been thought [85.6]. Accurate photoabsorption,

photodissociation and photoionization cross sections as

a function of wavelength for all the relevant species

are important parameters determining the reliability and

ultimate accuracy of an atmospheric model. Compi-

lations of data, such as that by Conway [85.7]and

Kirbyet al. [85.8] are very useful, but can become rapidly

outdated. The Smithsonian Astrophysical Observatory

maintains the world standard database, HITRAN, for

molecular line parameters and absorption cross sections

from the microwave through the ultraviolet for analysis

of atmospheric spectra [85.9]. Discussions of the needs

for atomic and molecular data in the context of space

astronomy, but including applications to atmospheric

physics, can be found in a book edited by Smith and

Wiese [85.10].

85.3 Tropospheric Warming/Upper Atmosphere Cooling

85.3.1 Incoming and Outgoing

Energy Fluxes

The overall temperature of a planet is determined by

a balance between incoming and outgoing energy ﬂuxes.

In a steady state, the planet must radiate as much en-

ergy as it absorbs from the sun. The Earth, radiating as

a black-body at an effective temperature T

, obeys the

Stefan–Boltzmann law in which the energy emitted is

expressed as σT

4πR

, with σ the Stefan–Boltzmann

Part G 85.3

1296 Part G Applications

constant, and R

the radius of the earth. An equation

expressing the equality of energy absorbed and energy

emitted can be written [85.11]as

πR

(1 − A) = σT

4πR

, (85.3)

where A is the albedo of the earth (the fraction of solar

radiation reﬂected from, rather than absorbed by, the

Earth), F

is the solar ﬂux at the edge of the earth’s

atmosphere, and πR

is the Earth’s area normal to the

solar ﬂux. Solving this equation for T

, one obtains

255 K (−18

◦

C).

The sun, which has a surface temperature of ap-

proximately 6000 K, emits most of its radiation in the

0.2–4.0 µm region of the spectrum (200–4000 nm). The

upper atmosphere of the Earth (thermosphere, iono-

sphere, mesosphere, and stratosphere) absorbs all the

solar radiation shortward of 320 nm. The atmosphere of

the earth absorbs only weakly in the visible region of

the spectrum where the solar ﬂux peaks.

The Earth, with an effective radiating temperature

of 255 K, emits mainly long-wavelength radiation in the

4–100 µm region. Molecules naturally present in the

atmosphere in trace amounts, such as carbon dioxide,

water and methane, absorb strongly in this wavelength

region [85.12]. Radiation coming from the Earth is

thus absorbed, reradiated back to the surface, and ther-

malized through collisions with the ambient gas. This

trapping of the radiation produces an additional warm-

ing of 33 K. Thus the mean surface temperature of

the Earth is 288 K, not 255 K as found for T

above.

This effect of the Earth’s atmosphere is known as the

greenhouse effect. The greenhouse effect is what makes

Earth habitable for life as we know it. Gases, both

natural and man-made, which absorb strongly in the

4–100 µm region, are known collectively as greenhouse

gases.

85.3.2 Tropospheric “Global” Warming

According to a 2000 National Research Council Report,

“the global-mean temperature at the earth’s surface is es-

timatedtohaverisenby0.25to0.4

◦

C during the past 20

years” [85.13]. The Intergovernmental Panel on Climate

Change (IPCC) has also concluded that global surface

temperatures have increased and that “there is new and

stronger evidence that most of the warming over the last

50 years is attributable to human activities” [85.14]. The

Arctic region has warmed by an estimated 1

◦

Cinthe

past two decades, leading to substantial changes in the

cryosphere [85.15]. Antarctic sea ice was stable from

1840 to 1950, but has since declined sharply. Sea ice

extent shows a 20% decline since about 1950 [85.16,17].

From air bubbles trapped at different depths in po-

lar ice, it is possible to determine carbon dioxide and

methane concentrations several thousand years ago.

Over the last two hundred years, CO

levels have

increased by 20%, from 280 to 330 ppm. Over the

next century the total amount of CO

in the atmo-

sphere since 1900 is expected to double to as much

as 600 ppm [85.18]. This increase is due primarily to the

burning of fossil fuels.

Although methane is present at levels several orders

of magnitude less than CO

, it is increasing much more

rapidly. Methane concentrations have more than dou-

bled over the last two hundred years due to industrial

processes, fuels, and agriculture [85.18].

The man-made chloroﬂuorocarbons (CFCs), which

have been widely used as refrigerants and in indus-

try, have been increasing in the atmosphere at a rate

of over 5% per year since the 1970s. Only recently

has there been an indication that this trend is slowing

down [85.19].

Ozone, which is a primary component of chemical

smog, is a pollutant when it occurs in the troposphere

and an effective greenhouse gas. It has been increasing

worldwide also.

This buildup of CO

,CH

, CFCs and tropospheric

causes a problem. In much of the spectral region from

5–100 µm, there is 100% absorption of radiation by the

atmosphere – due mainly to naturally occurring water

vapor. There is, however, a region of rather weak ab-

sorption, from ≈ 7–15µm, known as the “atmospheric

window”. Increased concentrations of the greenhouse

gases strengthen the absorption in this region, tending to

“close” this window, thus increasing the infrared opac-

ity of the atmosphere. The increased opacity causes an

immediate decrease in the thermal radiation from the

planet-atmosphere system, forcing the temperature to

rise until the energy balance is restored [85.20].

It is difﬁcult to prove that the buildup of greenhouse

gases is the cause of the observed temperature rise. Other

possible causes include slight changes in solar activity

and irradiance, and changes in ocean currents, which

may have a profound effect on global temperature and

climate. These are areas of active research.

Given the increase in concentrations of greenhouse

gases that has occurred and is predicted to continue,

the change in radiative heating of the troposphere can

be calculated. Models generally predict an increase in

tropospheric temperatures ranging from 1.5 to 4.5

◦

upon doubling the CO

concentrations over the next cen-

Part G 85.3

Applications of Atomic and Molecular Physics to Global Change 85.3 Tropospheric Warming/Upper Atmosphere Cooling 1297

tury. The 3

◦

C range in temperature is due to the ways

that different models incorporate climate feedbacks. Cli-

mate feedbacks include water vapor, snow and sea ice,

and clouds. Rising temperatures increase the concentra-

tion of water vapor, which is itself a greenhouse gas,

producing further warming. Rising temperatures reduce

the extent of reﬂective snow and ice, thus reducing the

Earth’s albedo. This leads to increased absorption of

solar radiation, further increasing temperatures. Clouds

both contribute to the albedo, thereby reducing the so-

lar ﬂux reaching the Earth, and absorb infrared radiation

causing temperatures to rise. The modeling of clouds

and their radiative properties is very difﬁcult, and is one

of the largest sources of uncertainty in the climate mod-

els. Understanding the role that the ocean, with its giant

heat capacity, plays in global warming, and identifying

and quantifying the various interactions occurring at the

ocean-atmosphere interface, are vital areas of research

which will affect the size of the predicted temperature

increase. At present, there are few obvious opportuni-

ties for traditional atomic and molecular physics to play

a signiﬁcant part in global-warming research.

85.3.3 Upper Atmosphere Cooling

The buildup of CO

has an even greater effect on the

temperature in the upper atmosphere than on that in the

troposphere [85.21]. As discussed in the Sect. 85.2.2,

is a coolant in the stratosphere, mesosphere and

thermosphere, but as a greenhouse gas is involved in

heating the troposphere. The explanation for this re-

volves around the collision physics issue of quenching

versus radiating.

In the troposphere, CO

absorbs infrared radiation

coming from the Earth, exciting the ν

vibrational bend-

ing mode at 15 µm. The excited molecule can either

reradiate or collisionally de-excite. In the lower atmo-

sphere where densities are large, the lifetime against

collisions is very short and the excited molecule is

rapidly quenched. This transfer of energy from radia-

tion through collisions into the kinetic energies of the

colliding partners results in a net heating.

In the stratosphere and above, atomic oxygen colli-

sions with CO

excite this same bending mode. But at

these higher altitudes, densities are lower and quench-

ing is greatly reduced. The excited molecule radiates

and the radiation escapes to space. A net cooling results

because the opacity is low at these altitudes.

Roble and co-workers [85.22, 23] have investigated

the doubling of CO

and CH

concentrations (as pre-

dicted for the next century) in the mesosphere and at

the lower boundary of the thermosphere. Using sophis-

ticated atmospheric general circulation models, they

predict that the stratosphere, mesosphere and thermo-

sphere will show signiﬁcant cooling — the largest

cooling of 40–50

◦

C occurring in the thermosphere.

The extent of this cooling very much depends

on the rate coefﬁcient for the O + CO

ex-

citation of the ν

bending mode. Rishbeth and

Roble [85.22] assumed a value for this rate coefﬁ-

cient of 1 × 10

−12

/s, intermediate between the value

of Sharma and Wintersteiner [85.24](6×10

−12

/s)

and an earlier value of 2 × 10

−13

/susedbyDick-

enson [85.25]. The Sharma and Wintersteiner value,

based on observations of 15 µm emission in the atmo-

sphere around 100–150 km, was recently conﬁrmed by

Rodgers et al. [85.26], but Pollock et al. [85.27] obtain

a value of 1.2×10

−12

/s in laboratory experiments.

Using the larger rate coefﬁcient would result in even

greater cooling [85.28].

The overall consequences of such a large tem-

perature decrease in the upper atmosphere have not

been fully explored — particularly the question as to

how the dynamics of the atmosphere will be affected.

Since many chemical reactions depend on temperature,

there may be considerable readjustments in the ver-

tical distribution of minor species in the atmosphere.

Cooler temperatures cause the atmosphere to contract,

reducing densities and, consequently, satellite drag.

Cooler temperatures may also increase the occurrence

of polar stratospheric clouds, thereby affecting ozone

depletion (Sect. 85.4).

Most signiﬁcantly, tropospheric warming and upper

atmosphere cooling both result from a buildup of CO

The size of the predicted cooling is greater by an order

of magnitude than the amount of the predicted heating.

Thus it may be possible to monitor the global warming

trend by observing the predicted cooling in the upper

atmosphere.

There is evidence in the mesosphere that this cool-

ing has already begun. Temperatures appear to have

decreased by 3–4

◦

C over the last decade [85.29, 30].

Gadsden[85.31] has also found that the frequency of oc-

currence of noctilucent clouds, the highest-lying clouds

in the atmosphere, has more than doubled over the last

twenty-ﬁve years. He has calculated that this change

could result from a decrease in the mean temperature

at the mesopause of 6.4

◦

C during this time period.

However, increased concentrations of water produced

by oxidation of increased amounts of methane may be

responsible for the more frequent appearance of the

clouds. This is an ongoing area of research.

Part G 85.3

1298 Part G Applications

85.4 Stratospheric Ozone

85.4.1 Production and Destruction

Ozone production takes place continually in the strato-

sphere during daylight hours, as molecular oxygen is

photodissociated and the resulting oxygen atoms un-

dergo three-body recombination with O

+ hν → 2O , (85.4)

O + O

+ M → O

+ M . (85.5)

Ozone can be destroyed through photodissociation:

+ hν → O

+ O , (85.6)

but because an oxygen atom is produced which imme-

diately recombines with another O

to form O

,nonet

loss of O

results. The photodissociation of O

and O

are important heating processes in the stratosphere.

The amount of ozone in the stratosphere is quite

variable, changing signiﬁcantly with the seasons and

with latitude. In the lower stratosphere, much of the

ozone is created over the equatorial regions and then

transported toward the poles.

Besides being photodissociated, O

is destroyed by

reactions with radicals that are involved in catalytic cy-

cles. The short-hand notation for the major cycles, NO

,ClO

(BrO

) refers to the catalytically active forms

involved in the cycles. Our knowledge about the relative

importance of these catalytic cycles in ozone destruc-

tion has increased dramatically over the last decade.

A number of these cycles are given below, with the

ozone-destroying step listed ﬁrst, and the rate-limiting

step closing the catalytic cycle and regenerating the

ozone-destroying radical, listed last. The net effect in

each of these cases is to convert ozone and atomic oxy-

gen (otherwise known as odd-oxygen) into molecular

oxygen:

NO + O

→ NO

+ O

+ O → NO + O

NET: O

+ O → 2O

, (85.7)

and

NO + O

→ NO

+ O

→ NO

+ O

+ hν → NO + O

NET: 2O

→ 3O

. (85.8)

OH + O

→ HO

+ O

→ OH + 2O

NET: 2O

→ 3O

, (85.9)

and the halogen cycle, in which Z = Cl or Br:

Z + O

→ ZO + O

ZO + O → Z + O

NET: O

+ O → 2O

. (85.10)

The following series of reactions couples the HO

and

halogen cycles:

+ ZO → HOZ + O

HOZ + hν → OH + Z

Z + O

→ ZO + O

OH + O

→ HO

+ O

NET: 2O

→ 3O

. (85.11)

Finally the reaction set

BrO + ClO → Br + Cl + O

Br + O

→ BrO + O

Cl + O

→ ClO + O

NET: 2O

→ 3O

(85.12)

is also important in the halogen destruction cycle. The

coupling between these different cycles by reactions

such as

+ NO → OH + NO

(85.13)

turns out to be very important in understanding the

details of ozone destruction, such as how much each

mechanism contributes to the destruction as a func-

tion of altitude and in the presence of aerosols.

Wennberg et al. [85.32] have recently shown that cat-

alytic destruction by NO

, which for two decades

was considered to be the predominant loss process,

accounted for less than 20% of the O

removal in

the lower stratosphere during May 1993. They further

show that the cycle involving the hydroxyl radical ac-

counted for nearly 50% of the total O

removal and

the halogen-radical chemistry was responsible for the

remaining 33%.

The NO

and HO

cycles are naturally occurring,

whereas the ClO

and BrO

cycles are due mainly to

man-made chemicals – the CFCs and halons. The am-

pliﬁcation that takes place through a catalytic cycle is the

Part G 85.4

Applications of Atomic and Molecular Physics to Global Change 85.4 Stratospheric Ozone 1299

reason that these chemicals, which are only present at

the level of parts per trillion, can have such a destructive

effect.

It is useful to think in terms of a total chemical bud-

get for a radical such as Cl which enters into a cycle.

Chlorine is put into the stratosphere when chemicals

such as CF

are released into the atmosphere. Such

compounds are chemically inert and insoluble in water,

and therefore are not easily cleansed out of the lower

atmosphere. In the stratosphere, however, the CF

is subjected to solar UV radiation and is photodisso-

ciated, producing the Cl radical. Chlorine exists in the

upper atmosphere in catalytically active forms, Cl and

ClO, as well as in stable, reservoir species, HCl and

ClONO

. The total chlorine budget consists of both

the catalytically active plus reservoir species. Reac-

tions which reduce the formation of reservoir species,

or convert reservoir species to catalytically active forms,

contribute to the ozone destruction. Photolysis of stable

reservoir species, such as ClONO

, can produce cat-

alytically active forms. Bromine has an identical cycle

to that of chlorine, but is 50 to 100 times more de-

structive than Cl because it does not react readily to go

into its reservoir form, HBr. A knowledge of the pho-

todestruction rates of all such species is important to

an understanding of the overall ozone photochemical

depletion problem.

Studies of the Antarctic ozone hole show that gas

phase photochemical cycles, as given above, are not

the whole story with respect to ozone depletion. Het-

erogeneous chemistries taking place on the surfaces

of ice crystals and sulfate aerosols play an important

role also. These are discussed brieﬂy in Sect. 85.4.2

and Sect. 85.4.4.

85.4.2 The Antarctic Ozone Hole

The ozone depletion problem was largely theoret-

ical until the discovery of the ozone hole over

Antarctica. Following the 1985 announcement by Far-

man et al. [85.33] of ground-based observations of

signiﬁcant decline in O

concentrations during spring-

time in the Southern Hemisphere, it was possible to map

this event using archived satellite data beginning in 1979.

The data depict a worsening event throughout the early

1980s. In 1987, 70% of the total O

column over Antarc-

tica was lost during the month of September and early

October, and the areal extent of the hole was ≈ 10% of

the Southern Hemisphere. The ozone hole has continued

to grow in depth and width [85.34]. Recent data shows

that this phenomenon continues, with the 2003 ozone

hole the second largest observed to date (the largest yet

observed was on September 10, 2000) [85.35].

The causal link between the release and buildup of

man-made CFCs and the ozone hole over Antarctica

has been quite convincingly established by Ander-

son et al. [85.36] through in situ observations from high

altitude aircraft ﬂights into the polar vortex during the

end of polar night and the beginning of Antarctic spring

in 1987.

The polar vortex is a stream of air circling Antarctica

in the winter, creating an isolated region which becomes

very cold during the polar night. Flights into the vortex

were able to document a heightened, increasing level of

ClO and a monotonically decreasing O

concentration

over a 3–4 week time period during late September and

early October.

The mechanism which appears to be repartitioning

the chlorine from its reservoir form into its catalyti-

cally active form is a heterogeneous process occurring

on the surfaces of polar stratospheric clouds. At the cold

temperatures during the polar night, polar stratospheric

clouds form, consisting of ice and nitric acid trihydrate.

Gaseous ClONO

collides with HCl that has been ad-

sorbed onto the surface of the cloud crystals. Chlorine

gas is liberated and the nitric acid formed in the reaction

remains in the ice [85.36]:

HCl + ClONO

→ Cl

(g) + HNO

. (85.14)

As solar radiation starts to penetrate the region at

the beginning of spring, the Cl

molecules are rapidly

photodissociated, producing Cl atoms which initiate the

catalytic destruction of O

As there are no oxygen atoms around to complete

the catalytic cycles, several mechanisms for regenerating

the Cl and Br radicals have been proposed which involve

only the ClO and BrO molecules themselves.

Mechanism I [85.37]

ClO + ClO + M → (ClO)

+ M

(ClO)

+ hν → Cl + ClOO

ClOO + M → Cl + O

+ M

2×(Cl + O

→ ClO + O

)

NET: 2O

→ 3O

; (85.15)

Mechanism II [85.38]

ClO + BrO → Cl + Br + O

Cl + O

→ ClO + O

Br + O

→ BrO + O

NET: 2O

→ 3O

. (85.16)

Part G 85.4

1300 Part G Applications

While Mechanism I accounts for 75% of the ob-

served ozone loss, the sum of I and II yields a destruction

rate in harmony with the observed O

loss rates [85.36].

85.4.3 Arctic Ozone Loss

The region around the North Pole does not appear

to exhibit an ozone hole as severe as that found in

the Antarctic. Several factors lessen the probability of

a signiﬁcant ozone hole developing in the Arctic. First,

a stable polar vortex does not get well established due to

increased atmospheric turbulence from the greater land

surface area in the Northern Hemisphere. Second, tem-

peratures during the Arctic winter do not get as cold as

during the Antarctic winter, so that polar stratospheric

clouds (PSCs) do not form as easily. As seen in the pre-

ceding section, the surfaces of PSCs play an essential

role in the O

destruction mechanisms in the Antarctic.

However, ozone levels are showing 20–25% reduc-

tions during February and March [85.39] over a much

larger area around the North Pole than in the South.

Thus ozone destruction is taking place during the tran-

sition from polar winter to spring in the Arctic but the

phenomenon is more widespread, diffuse, and not as

well-contained as in the Antarctic.

85.4.4 Global Ozone Depletion

Over the last twenty-ﬁve years, satellite instruments

have measured the total ozone column in the atmosphere.

During this time ozone levels have been steadily decreas-

ing globally, especially at mid- to high-latitudes. Recent

analysis indicates the ﬁrst evidence of recovery of strato-

spheric ozone levels, with diminished rates of ozone loss

at altitudes of 35–45 km, coupled with a slowdown in

the increase in stratospheric loading of chlorine [85.40].

Heterogeneous reactions on aerosol surfaces, as well

as the homogeneous gas phase chemical cycles men-

tioned earlier, must be invoked to explain the global

decline in ozone levels. A particularly important reac-

tion appears to be the hydrolysis of N

on sulfate

aerosols. This occurs very rapidly, converting reactive

nitrogen, NO

, into its reservoir species HNO

+ H

sulfate aerosol

−−−−−−−−−→ 2HNO

. (85.17)

The N

is formed at night by reaction of NO

and NO

. Following the hydrolysis of N

,there

is less reactive NO

around to convert ClO into its

reservoir species, ClONO

,andlessNO

around to

convert OH into the reservoir species, HNO

. A height-

ened sensitivity of the ozone to increasing levels of

CFCs develops [85.41]. It has been shown that cer-

tain regions of signiﬁcantly depleted ozone also show

high concentrations of sulfate aerosols. In addition,

measurements of the ratio of catalytically active ni-

trogen to total nitrogen can be reproduced using the

above heterogeneous reaction, and not by using gas

phase processes alone. Study of further mechanisms

at varying altitudes and latitudes is an active area of

research.

Record low global ozone measurements, 2% to 3%

lower than any previous year, were reported beginning

in 1992 [85.42] and continuing well into 1993. The in-

crease in naturally occurring aerosols due to the eruption

of Mount Pinatubo in June 1991 appears to explain this

decline. During the winter of 1993–1994, total ozone

levels returned to levels slightly above normal [85.43],

presumably because the excess aerosols had been re-

moved from the stratosphere by natural sedimentation

processes.

The continuing buildup of CO

is predicted to

contribute to increased cooling of the stratosphere. De-

clining temperatures in the stratosphere may increase

the frequency of formation of polar stratospheric clouds

which drive the destructive heterogeneous chemistry

creating the Antarctic ozone hole. An increased occur-

rence of these clouds outside of the polar regions could

affect ozone levels globally. There are also indications

that certain ozone depletion chemistries taking place on

the surface of sulfate aerosols may also be enhanced by

lower temperatures [85.41].

Ozone itself is the dominant heat source in the lower

stratosphere. Decreasing the amount of ozone drives

temperatures still lower [85.44]. It is unfortunate that the

two most signiﬁcant atmospheric global change effects

— the buildup of CO

and the enhanced ozone destruc-

tion due to man-made CFCs — both cause decreasing

temperatures in the stratosphere which may further en-

hance the destructiveness of the ozone photochemical

cycles.

85.5 Atmospheric Measurements

Ground-based observations, as well as measurements

made by instruments carried aloft in satellites, bal-

loons, and high-ﬂying aircraft, allow one to explore the

atmosphere.

Part G 85.5

Applications of Atomic and Molecular Physics to Global Change References 1301

Measurements may be made either in situ or by

remote-sensing techniques. The region of the atmo-

sphere from ≈ 60 km to 120 km, encompassing the

mesosphere and lower thermosphere and ionosphere,

cannot be studied in situ as it is too high for balloons

and aircraft, and too low (i. e., too much drag) for satel-

lites. For this region, remote sensing experiments are

essential and a comprehensive book on the subject is

recommended [85.45].

Most of the instruments used to make atmo-

spheric measurements have been developed in molecular

physics and spectroscopy laboratories. Even experi-

ments utilizing sophisticated techniques, such as laser

induced ﬂuorescence, and instruments, such as Fourier

transform spectrometers, are being ﬂown on payloads.

An excellent compendium of ozone-measuring instru-

ments for stratospheric research has been assembled by

Grant [85.46].

A combination of good laboratory experiments, the-

oretical calculations, and ingenuity are necessary to

extract accurate information from measurements made

in the atmosphere. For instance, in order to understand

the complicated interactions of the different photochem-

ical cycles involved in ozone chemistry, spectroscopic

emissions and absorptions of the many trace species

are used to measure concentration proﬁles. An accu-

rate knowledge of the emission spectroscopy of species

such as OH, HO

O, O

, HNO

,NO

, HNO

,ClNO

, BrO, HCl, HOCl, and ClO is

essential. Such measurements provide a rigorous test of

atmospheric models. The recently-launched NASA EOS

Aura satellite carries instruments that will make global

measurements of a number of these species [85.47].

In order to analyze the data and deconvolve some

of the line proﬁles to give information on concentra-

tions as a function of altitude, molecular data such as

line strengths and pressure broadening coefﬁcients are

needed [85.9].

Until recently, it has been impossible for remote-

sensing experiments to distinguish between ozone

occurring in the stratosphere (where it is formed nat-

urally) and ozone occurring in the troposphere (where it

is a pollutant). Satellite instruments such as the ESA

Global Ozone Monitoring Experiment (GOME), the

Scanning Imaging Absorption Spectrometer for Atmo-

spheric Chartography (SCIAMACHY) and the Ozone

Monitoring Instrument (OMI) have broad enough spec-

tral coverage and high enough resolution that the

temperature dependence of the ozone absorption fea-

tures from 300–340 nm, known as the Huggins bands,

can be used to separate out the ozone concentrations in

the middle and lower atmospheres [85.48, 49].

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Part G 85