Salby M.L. Fundamentals of Atmospheric Physics

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1 A Global

View

Geological evidence suggests that primitive forms of plant life developed

aqua

deep in the oceans, at a time when the earth's atmosphere contained little

or no oxygen and damaging UV radiation passed freely to the planet's surface.

Through photosynthesis, these early forms of life are thought to have liberated

oxygen, which then passed to the surface where it was photodissociated by UV

radiation according to the reaction

0 2 -+-

hv --+

20. (1.26)

Atomic oxygen produced by (1.26) could then recombine with 02 to form

ozone in the termolecular reaction

02 + 0 +M--+ 03 +M,

(1.27)

where M represents a third body needed to carry off excess energy liberated

by the combination of O and 02. Ozone created in (1.27) is dissociated by

UV radiation according to the reaction

0 3 +

hv-+

02 +0.

(1.28)

If third bodies are abundant, atomic oxygen produced by (1.28) recombines

almost immediately with 02 in (1.27) to again form ozone. Thus, reactions

(1.27) and (1.28) constitute a "closed cycle" that involves no net loss of com-

ponents. Since the only result is the absorption of solar energy, this cycle can

process UV radiation very efficiently. By removing harmful UV from the so-

lar spectrum, ozone is thought to have allowed life to spread upward to the

oceans' surfaces, where it had greater access to visible radiation, could pro-

duce more oxygen through photosynthesis, and was then able to evolve into

more sophisticated forms.

The zonal-mean distribution of ozone mixing ratio is shown in Fig. 1.17 as a

function of latitude and altitude. Whereas atmospheric water vapor is confined

chiefly to the troposphere, ozone is concentrated in the stratosphere. Ozone

mixing ratio increases sharply above the tropopause, reaching a maximum of

about 10 ppmv near 30 km (10 mb). The zonal-mean ozone mixing ratio ?o3

is largest in the tropics, where the flux of solar UV and photodissociation of

02 are large.

The photochemical lifetime of ozone varies sharply with altitude. In the

lower stratosphere, ozone has a photochemical lifetime of several weeks. Since

this is long compared to the characteristic timescale of air motion (,~ 1 day),

ro3 behaves as a tracer at these altitudes and its distribution is controlled by

dynamical influences. Should ozone find its way into the troposphere, it is

quickly destroyed. Its water solubility makes 03 readily absorbed by convec-

tive systems, which precipitate it to the surface, where it can be destroyed by a

variety of oxidation processes. Thus, the troposphere serves as a sink of strato-

spheric ozone. The photochemical lifetime of ozone also decreases upward, to

of order 1 day by 30 km and only 1 h by the stratopause. For this reason, the

1.2 Composition and Structure

0.05 ),

0.1

0.2

0.4

0.5

0.7

.13

E 1.0

1.5

a::: 2.0

03 3.0

w 5.0

n..' 7.0

13..

I00

-90

OZONE (ppmv)

-60 -30 0 30 60

LATITUDE

Figure 1.17 Zonal-mean mixing ratio of ozone (contoured) and density of ozone (shaded)

averaged over January-February 1979, as functions of latitude and pressure, obtained from the

Limb Infrared Monitor of the Stratosphere (LIMS) on board Nimbus-7. The shaded levels cor-

respond to

20, 40, and 60% of the maximum value.

distribution of ozone in the upper stratosphere and mesosphere is controlled

chiefly by photochemical influences.

Even though its mixing ratio ?o3 maximizes near 30 km, atmospheric ozone

is concentrated in the lower stratosphere (Fig. 1.17). Because air density de-

creases exponentially with altitude, the density of ozone Po3 (shaded area) is

concentrated at altitudes of 10 to 20 km (1.24). Largest values are found in

a shallow layer near 30 mb in the tropics, which descends and deepens in ex-

tratropical regions. The column abundance, or

total ozone,

Eo3 is expressed in

Dobson units (DU), which measure in thousandths of a centimeter the depth

the ozone column would assume if brought to standard temperature and pres-

sure. Figure 1.18 shows the zonally averaged column abundance of ozone as a

function of latitude and season. Values of ~o3 range from about 250 DU near

the equator to in excess of 400 DU at high latitudes. The entire ozone column

measures less than one-half of 1 cm at standard temperature and pressure!

Even though most stratospheric ozone is produced in the tropics, the greatest

column abundances are found at middle and high latitudes. Like the temper-

ature distribution in the mesosphere (Fig. 1.7), this peculiarity of the ozone

distribution illustrates the importance of dynamics to observed composition

and structure.

As is true for water vapor, the circulation plays a key role in determining

the mean distribution of 03 and makes the global distribution on individ-

ual days dynamic. Figure 1.19 shows the distribution of ~o3 over the Northern

32 1 A Global

View

OZONE COLUMN ABUNDANCE (DU)

K',,,~:~o6~--'--):--)---~.---," a'zo. -

u.i 3o

v-- 0

F--

..j -:30

-60

-90

d F M A M d d A S 0 N O

MONTH

Figure

1.18 Zonal-mean column abundance of ozone, or

total ozone,

as a function of latitude

and month. Based on the historical record prior to 1980. From London (1980).

and Southern Hemispheres on individual days, as observed by the Total Ozone

Mapping Spectrometer (TOMS). The distribution over the Northern Hemi-

sphere (Fig. 1.19a) has been arranged by the circulation into several anomalies

in which ~o3 varies by as much as 100%. Figure 1.19a, which is contemporane-

ous with the 500-mb circulation in Fig. 1.9a, reveals a strong correspondence

between variations of ~o3 and synoptic weather systems in the troposphere.

Both evolve on a timescale of a day. The distribution over the Southern Hemi-

sphere (Fig. 1.19b) is distinguished by column abundances of less than 200 DU

(white) that delineate the '~ntarctic ozone hole." Anomalously low ~o3 ap-

pears over the South Pole each year during Austral spring and then disappears

some two months later. The formation of the ozone hole, which emerged in

the 1980s, is attributed to increasing levels of atmospheric chlorine. Its disap-

pearance each year occurs through dynamics.

METHANE

Several other trace gases also figure importantly in radiative and chemi-

cal processes. These include species that are produced naturally, like methane

(CH4) ,

and species that are produced solely by human activities, like chlorofiu-

orocarbons (CFCs). Methane is produced primarily by bacterial and surface

processes that occur naturally. However, anthropogenic sources such as min-

ing and industrial activities may constitute as much as 20% of

CH 4

production.

Figure

1.19 Distribution of total ozone (color) from the Total Ozone Mapping Spectrom-

eter (TOMS) on board Nimbus-7 and pressure (contours) on the 375 K isentropic surface (see

Sec. 2.4.1) over (a) the Northern Hemisphere on March 4, 1984 and (b) the Southern Hemisphere

on October 25, 1983.

1.2 Composition and Structure

8o[

I-

.._1

< C-1

10-11 10-1o

N2 t "~ i

" i -

',,,,,, "",, icck

i -

,,.,1.

i i

r i

i I; i I ii i

10-9 10 8 10 -7 10 6 10 -5 10 4 10 -3

VOLUME MIXING RATIO

Figure 1.20

Mixing ratios of radiatively active trace species as functions of altitude.

Source:

Goody and Yung (1989).

Methane is long lived and therefore well mixed in the troposphere, where it

has a constant mixing ratio of rCH" "~ 1.7 ppmv (Fig. 1.20). In the stratosphere,

rCH" decreases with altitude as a result of oxidation. This process ultimately

leads to the formation of stratospheric water vapor and is thought to be re-

sponsible for the increase of ru2 o observed in the stratosphere.

Like CO2, methane concentrations are increasing steadily. Although the

exact rate is a matter of some debate, an increase of 1 to 2% per year is

generally agreed on. Increasing levels of

CH 4 are

predicted by GCM studies to

lead to increased global temperature, but not as great as those anticipated from

increases of CO2. Methane is also involved in the complex photochemistry of

ozone.

CHLOROFLUOROCARBONS

Industrial halocarbons like

CC14

(CFC-10) and

CFC13

(CFC-11) and

CF2C12

(CFC-12) are used widely as aerosol propellants, in refrigeration, and in a va-

riety of manufacturing processes. As shown in Fig. 1.21, the release of these

gases into the atmosphere has increased steadily since the Second World War,

although the rate of increase has abated in the wake of heightened environ-

mental concern and reduced demand. These anthropogenic species are stable

in the troposphere, where their water insolubility makes them immune to nor-

mal scavenging processes associated with precipitation. Because they are long

lived, CFCs are well mixed in the troposphere, CFC-11 and CFC-12 having

nearly uniform mixing ratios of rcFc_11 = 0.2 ppbv and rCFC_12

0.3 ppbv.

Their radiative properties involve industrial halocarbons in the energy balance

of the earth, making them another consideration for climate change.

1 A Global View

106F I

I I I I I I I I

9 ATMOSPHERIC RELEASE

OF CHLOROFLUOROCARBONS

jj~

105 - s'~

9

i J

!C-10

w I

CO ~t~ 4 I

<t:: .u - I

I,.iJ 4'

...,J

LI.,I I

13: I

,..,j !

< O

::3 I

z CFC- s

z 12 CFC-11

< 03 I

1 - t

I '

2 I

1940 1950 1960 1970 1980

TIME (Year)

Figure 1.21 Atmospheric release of CFCs. After Brasseur and Solomon (1986).

The greatest interest in CFCs surrounds their impact on the ozone layer. Be-

cause they are long lived in the troposphere, CFCs are eventually transported

into the stratosphere, where UV radiation photodissociates them. Although

constant in the troposphere, mixing ratios of CFC-11 and -12 decrease in the

stratosphere (Fig. 1.20). Free chlorine liberated via dissociation of CFCs can

destroy ozone. Catalytic destruction of 03 by C1 is responsible for the forma-

tion of the Antarctic ozone hole in Fig. 1.19b. Chlorine is but one piece of

the scientific puzzle surrounding the ozone hole. Very high clouds that rarely

form except over Antarctica also play a key role, as does the stratospheric

circulation during the spring breakdown of the ozone hole.

NITROGEN COMPOUNDS

Oxides of nitrogen, such as nitrous oxide

(N20)

and nitric oxide (NO),

are also relevant to the photochemistry of ozone. Nitrous oxide is produced

primarily by natural means relating to bacterial processes in soils. Anthro-

pogenic sources of N20 include nitrogen fertilizers and combustion of fossil

1.2 Composition and Structure

fuels, which may account for as much as 25% of the total production. These

sources have altered the natural cycle of nitrogen by introducing it in the form

of NeO instead of N e. Like methane, NeO is long lived and therefore well

mixed in the troposphere, with a nearly uniform mixing ratio of rN2 o = 300

ppbv.

In the stratosphere, rN:o decreases with altitude due to dissociation of

nitrous oxide, which represents the primary source of stratospheric NO. Like

free chlorine, NO can destroy ozone catalytically. Nitric oxide is also produced

as a by-product of inefficient combustion, for example, in aircraft exhaust.

Nitrous oxide, similar to other anthropogenic gases, has increased steadily

in recent years. Beyond its relevance to ozone, the increasing level of NaO

has implications to the energy budget of the earth because nitrous oxide is

radiatively active, albeit much less so than CO2.

ATMOSPHERIC AEROSOL

Suspensions of liquid and solid particles are relevant to radiative as well as

chemical processes. Ranging in size from thousandths of a micron to several

hundred microns, aerosol particles are vital to atmospheric behavior because

they promote cloud formation. Aerosol particles serve as condensation nuclei

for water droplets and ice crystals, which do not form readily in their absence.

These small particulates are produced naturally, for example, as dust, sea salt,

and volcanic debris. They are also produced anthropogenically through com-

bustion and industrial processes. An important source of aerosols is

gas-to-

particle conversion,

which occurs through chemical reactions involving, among

other precursors, the gaseous emission sulfur dioxide (SOe) that is produced by

industry and naturally by volcanos. Reflecting anthropogenic sources, aerosol

concentrations are high in urban areas and near industrial complexes, for ex-

ample, as much as an order of magnitude greater than concentrations over

maritime regions (Fig. 1.22). Even higher concentrations are found in as-

sociation with windblown silicates during desert dust storms. In addition to

their role in cloud formation, aerosol particles scatter solar radiation at visi-

ble wavelengths and absorb IR radiation emitted by the planet's surface and

atmosphere. These radiative properties involve aerosols in the energy balance

of the earth, wherein their influence is thought to be as great as that of COe.

Aerosols play key roles in several chemical processes. One of the more

notable is heterogeneous chemistry, which requires the presence of multiple

phases. Reactions involving chlorine, which take place on the surfaces of cloud

particles, lie at the heart of the chain of events that culminates in the forma-

tion of the Antarctic ozone hole in Fig. 1.19b. Beyond this role, aerosols are

thought to figure in climatic fluctuations and possibly in the evolution of the

atmosphere. Sporadic increases of aerosol produced by major volcanic erup-

tions have been linked to changes in thermal, optical, and chemical properties

of the atmosphere.

36 1 A Global View

100C

I I I I I ! I

900

Dust~

Storm ~

800~ ,'

"<' 700

r ,

~o 600 , ',

~~ : :

g ~ 500[ '-,

~ v u

i~) (/'}L'~ O r / I Urban ,

, -

9 "o

40 :

8 '

,,=,

< aoo~- ',,, -

2001-

II --

100[ ',,

-' "-

10-3 10-2 10-1 10 0 101 10 2 10 a

RADIUS a (pm)

Figure 1.22 Size spectrum of aerosol surface area, as a function of particle radius. The

representation is area-preserving: Equal areas under the curves represent equal number densities

of aerosol particles.

Source:

Slinn (1975).

1.2.5 Clouds

One of the most striking features of the earth, when viewed from space, is its

extensive coverage by clouds. At any instant, about half of the planet is cloud

covered. Clouds appear with a wide range of shapes, sizes, and microphysi-

cal properties. Like trace species, the cloud field is highly dynamic because

of its close relationship to the circulation. Beyond these more obvious char-

acteristics, clouds and related phenomena figure importantly in a variety of

atmospheric processes.

The reason clouds are so striking from space is that they reflect back to

space a large fraction of incident solar radiation, which is concentrated at

visible wavelengths. Owing to their high reflectance at visible wavelengths,

clouds shield the planet from solar radiation, which is one important role they

play in the earth's energy budget. Clouds also figure importantly in the budget

of terrestrial radiation that is emitted in the IR by the planet's surface and

atmosphere to offset solar heating. Because they absorb strongly in the IR,

cloud particles of water and ice increase the atmosphere's opacity to terrestrial

radiation and hence its trapping of radiant energynear the earth's surface.

Their optical properties allow clouds to be observed from space in measure-

ments of visible and IR radiation. Figure 1.23 shows an image from the l l-/zm

1.2 Composition and Structure

~ii~iiiii~i~!~ 84184 .....

Figure

1.23 Infrared image at 1200 GMT on March 4, 1984, from the ll-/xm channel of

Meteosat-2. Gray scale displays equivalent blackbody temperature from warmest (black) to coldest

(white). (See Fig. 1.24 for geographical landmarks.) Supplied by the European Space Agency.

channel of Meteosat-2 that is contemporaneous with the water vapor image in

Fig. 1.15. The gray scale in Fig. 1.23 emphasizes the highest objects emitting

in the IR. Bright areas correspond to cold high clouds, whereas dark areas

indicate warm surfaces under cloud-free conditions. Many features in the IR

image are well correlated with features in the water vapor image. This is es-

pecially true of high convective clouds in the tropics, which extend upward to

the tropopause. Having horizontal dimensions of tens to a few hundred kilo-

meters, those clouds mark deep convective towers that displace surface air

vertically on timescales of hours to a day. On the other hand, the IR image

reveals the subtropical Sahara and Kalahari deserts in dark areas that are far

less evident in the water vapor image.

High clouds at midlatitudes also correspond well with features in the water

vapor image. Most striking is the band of high cirrus that extends northeast-

ward from South America (Fig. 1.24), across the Atlantic, and into Africa.

Overlying the tongue of moisture in Fig. 1.15, this high cloud cover is part of

1 A Global

View

Figure

1.24 Visible image at 1200 GMT on March 4, 1984, from Meteosat-2. Supplied by

the European Space Agency.

a frontal system that accompanies the cyclone in the eastern Atlantic, which

appears in Fig. 1.23 as a spiral of lower clouds (compare Fig. 1.9a). As it ad-

vances, that cyclone draws moist tropical air poleward ahead of the front and

entrains it with drier midlatitude air.

Clouds also appear in visible imagery, like that shown in Fig. 1.24, which

is contemporaneous with the IR and water vapor images. Unlike those im-

ages, which follow from emission of IR radiation, the visible image represents

reflected solar radiation, so it does not depend on the temperature of the ob-

jects involved. Consequently, low and comparatively warm stratiform clouds

appear just as prominently as high cold convective clouds. The prevalence of

stratiform features in Fig. 1.24 indicates that even shallow clouds are efficient

reflectors of solar radiation. Their extensive coverage of the globe makes such

clouds an important consideration in the earth's radiation budget. Swirls of

light and dark in the water vapor image, which are signatures of horizontal

motion, are only weakly evident in the IR image and are completely absent

from the visible image.

1.2 Composition and Structure

The roles clouds play in the budgets of solar and terrestrial radiation make

them a key ingredient of climate. In fact, the influence cloud cover exerts

on the earth's energy balance is an order of magnitude greater than that

of CO2. Its dependence on the circulation, thermal structure, and moisture

distribution, make the cloud field an especially interactive component of the

earth-atmosphere system.

With the exception of shallow stratus, most clouds in Figs. 1.23 and 1.24

develop through vertical motion. Two forms of convection are distinguished in

the atmosphere. Cumulus convection, which is often implied by the term con-

vection alone, involves thermally driven circulations that operate on horizontal

dimensions of order 100 km and smaller. Deep tropical clouds in Fig. 1.23 are

a signature of cumulus convection, which displaces surface air vertically on

small horizontal dimensions. Sloping convection is associated with forced lift-

ing, when one body of air overrides another, and occurs coherently over large

horizontal dimensions. The band of high cloud cover preceding the cyclone in

the eastern Atlantic is a signature of sloping convection.

Beyond its involvement in radiative processes, convection plays a key role

in the dynamics of the atmosphere and in its interaction with the oceans.

Deep convection in the tropics liberates large quantities of latent heat that

are released when water vapor condenses and precipitates back to the surface.

Derived from heat exchange with the oceans, latent heating in the tropics

represents a major source of energy for the atmosphere. It is for this reason

that deep cumulus clouds are used as a proxy for atmospheric heating.

Figure 1.25a shows a nearly instantaneous image of the global cloud field, as

constructed from 11-/zm radiances measured aboard six satellites. The highest

(brightest) clouds are found in the tropics in a narrow band of cumulus con-

vection. Known as the Inter Tropical Convergence Zone (ITCZ), this band of

organized convection is oriented parallel to the equator, except over the trop-

ical landmasses: South America, Africa, and the "maritime continent" over

Indonesia, where the zone of convection widens. Inside the ITCZ, deep con-

vection is supported by the release of latent heat when moisture condenses.

Transfers of moisture and energy make the ITCZ important to the tropical

circulation and to interactions between the atmosphere and oceans.

The ITCZ emerges prominently in the time-mean cloud field, shown in

Fig. 1.25b. Over maritime regions, the time-mean ITCZ appears as a nar-

row strip parallel to the equator, which reflects the convergence of surface air

from the two hemispheres inside the Hadley circulation. Over tropical land-

masses, time-mean cloud cover expands due to the additional influence of

surface heating, which triggers convection diurnally. There, as throughout the

tropics, the unsteady component of the cloud field is as large as the time-

mean component. In addition to operating on timescales of hours to a day,

organized convection migrates north and south annually with the sun and in-

volves the monsoons over southeast Asia and northern Australia during the

solstices.