Space Sci Rev (2012) 168:1–21DOI 10.1007/s11214-011-9791-y

Interactions Between the Lower, Middle and UpperAtmosphere

Anne K. Smith

Received: 2 February 2011 / Accepted: 30 May 2011 / Published online: 19 June 2011© Springer Science+Business Media B.V. 2011

Abstract This survey of atmospheric regions is based on a presentation given at the Work-shop on Coupling between Earth’s Atmosphere and its Plasma Environment. The workshopwas sponsored by the International Space Science Institute (ISSI) during September 2010.This survey begins with descriptions of aspects of the troposphere, stratosphere, and meso-sphere that affect vertical coupling. Then the tropopause, stratopause, and mesopause areconsidered, with a discussion of how they inhibit or allow the transmission of atmosphericmass or perturbations. Two transition interfaces are emphasized: the tropopause, which sep-arates the troposphere from the middle atmosphere, and the mesopause, which separates themiddle atmosphere from the thermosphere. The survey of coupling processes looks at ra-diative processes, wave propagation, and mass exchange. Each topic includes one or morespecific examples drawn from recent literature.

Keywords Coupling · Solar variability

1 Introduction

This article introduces the vertical structure of the atmosphere and discusses the interfacesbetween regions or layers. The focus is on the interactions across regions and the impedi-ments to interaction. Brief descriptions of the lower and middle atmosphere are given andthe regions are contrasted. Several examples of coupling across regions are presented. Theadditional articles in this volume include detailed discussions of aspects of these interactionsand other processes not discussed here.

2 Definition of the Atmospheric Regions

Atmospheric scientists classify the neutral atmosphere in regions. The regions and the tran-sitions between them are identified by their mean thermal structure. They have traditionally

A.K. Smith (�)NCAR, Boulder, CO, USAe-mail: aksmith@ucar.edu

2 A.K. Smith

been considered to be layers but, since the atmosphere is not layered globally, the termregions will be used here. The regions are the troposphere (Earth’s surface to approxi-mately 10–18 km), the stratosphere (10–18 km to about 50 km), the mesosphere (about50 to 100 km) and the thermosphere (above 100 km). The transitions between these regionsare the tropopause, the stratopause, and the mesopause. Because of common dynamicaland chemical environments, the region that encompasses the stratosphere, mesosphere andsometimes the lowermost part of the thermosphere is also called the middle atmosphere.

2.1 The Troposphere

The troposphere is strongly affected by the Earth’s surface properties.

– The orography of the surface affects the generation of atmospheric waves as winds flow-ing near the surface are displaced horizontally or vertically.

– The landforms, vegetation, and snow cover of land surfaces affect absorption versus re-flection of incoming solar radiation and the emission of outgoing infrared radiation.

– The landforms, vegetation, and snow cover of land surfaces affect exchange of water andtrace chemical species with the atmosphere.

– The surface properties (for example, surface wave amplitudes) and ice cover of ocean andlake surfaces affect absorption versus reflection of solar radiation and the exchange ofwater and trace chemical species with the atmosphere.

The troposphere is highly variable on short time scales; the details of this variability areoften difficult to predict. Basic physical processes act to create unstable conditions. Theresponses to the unstable conditions are a variety of dynamical phenomena that act to bringthe atmosphere back toward stability.

An important mode of instability is convective instability. The bulk of the solar flux thatenters the troposphere passes through the atmosphere and is absorbed at the Earth’s surface(IPCC 2007). A substantial portion of the energy is converted to heat. An additional portionof the energy evaporates water that then contributes to the humidity of the atmosphere. Thetransfer of the heat from the land or ocean surface to the atmosphere occurs by radiative(infrared), sensible, or latent heating. Reflection of solar radiation at the surface or by par-ticles within the atmosphere also affects the energy exchange between the surface and theatmosphere.

The heat input to the lower levels of the atmosphere can create convective instability. Indry convective instability (i.e., under conditions of low humidity), the air will be unstablewherever the temperature structure satisfies

∂T

∂z< − g

cp

≡ �d (1)

where T is temperature, g is the acceleration of gravity and cp in the heat capacity of air atconstant pressure. The so-called “dry adiabatic lapse rate” �d described by (1) is −9.8 K/km.

If dry air is displaced upward, it will cool at the rate of 9.8 K/km. By definition, if dryair is convectively unstable, its temperature falls off more rapidly than that. As a result,air displaced upward will be warmer, and therefore lighter, than the surrounding air. Thedisplaced air feels an upward force rather than a restoring force; it can continue to riseuntil it reaches an altitude where the temperature profile is no longer convectively unstable.This rising motion is convection. The opposite condition is an atmospheric profile that isconvectively stable. In that case, air displaced upwards will cool and become colder than the

Interactions Between the Lower, Middle and Upper Atmosphere 3

surrounding air. Gravitational force will then be a restoring force to accelerate the displacedair back towards its original altitude.

When the humidity is high enough, the exchange of heat due to the condensation ofmoisture has an impact on the convection process. Air that ascends and cools can hold lessmoisture by virtue of its cooler temperature. As the relative humidity becomes high enoughfor the moisture to condense into cloud droplets, latent heat will be released. The latent heatrelease slows the rate at which the air cools. Under saturated conditions the moist adiabaticlapse rate �s is

�s = �d − L

cp

dws

dz(2)

where ws is the mixing ratio at which water becomes saturated and L is the latent heat ofcondensation. The moist adiabatic lapse rate depends on both humidity and temperature.The lapse rate normally becomes steeper in the upper troposphere where both humidity andtemperature are lower. The lapse rate is also expected to become less steep as the overalltemperature increases due to global warming (Held and Soden 2006).

Due to conservation of mass, convective upwelling must be balanced by downwelling ofair in surrounding areas and the upwelling and downwelling air masses must be linked byhorizontal motion. The combined effect of the upward and downward mass flow keep themean atmosphere temperature structure over large parts of the Earth near the moist adiabaticlapse rate. The horizontal scale of convection is small (less than a few km) even though theheating that generates the instability can be widespread. The location and timing of theappearance of individual convective cells is difficult to predict. The atmospheric variablesthat best predict where convection will occur are discussed e.g. by Sherwood (1999).

Convective instability is one of several important modes of instability. It is most im-portant in the tropics where the heating is normally strong and the horizontal variations inangular momentum are the weakest. However, convection also occurs in midlatitudes, par-ticularly over land surfaces during the warmer times of the year.

The rotation of the Earth is associated with a change in the angular momentum fromthe equator (largest) to the poles (zero). The change in the angular momentum affects thehorizontal path of air displaced either vertically or in latitude. Instabilities in which theconservation of angular momentum play a role are barotropic and baroclinic instability.Baroclinic instability is a key component of the generation of midlatitude storm systems.

Baroclinic instability is an inevitable situation for the Earth’s atmosphere. The unstablesituation develops because the atmosphere is heated more strongly in the tropics and there-fore has an equator to pole temperature gradient. It also depends on the rotating Earth andthe importance of the Coriolis effect on horizontal motion. Through the thermal wind re-lation, the meridional temperature gradient is associated with a vertical shear in the zonalwind.

∂u

∂z≈ − R

Hf

∂T

∂y(3)

where u is the zonal wind, R is the gas constant for air, f is the Coriolis term, and H is theatmospheric scale height. Small north-south perturbations in the air flow advect cooler airin the equatorward direction and warmer air poleward. Depending on the spatial scale andfrequency of the perturbation and on the conditions of the background atmosphere, somedisturbances will be unstable and their amplitudes can grow exponentially. During winter,baroclinic instability leads to the steady generation of midlatitude storm systems.

With vigorous moist convection in the tropics, transport of air parcels from the bottomto the top of the troposphere can occur on timescales of minutes. The timescale for vertical

4 A.K. Smith

transport is also fairly short (a few days) in midlatitudes. As a result, all but water and short-lived chemical species in the troposphere are well-mixed in the vertical.

Due to instabilities, variations in external heating, and other processes, the winds, tem-perature, and humidity in the troposphere vary with space and time on multiple scales. Someof these variations damp out with time and have primarily a local impact. However, manyperturbations can propagate in the horizontal and/or vertical direction. Propagating distur-bances are atmospheric waves. Their behavior in space and time can be represented by awave equation in one or more dimensions. The generation and propagation of waves is adominant aspect of the atmosphere at all levels and is a key component linking differentregions.

Waves generated in the troposphere that have important impacts on the middle atmo-sphere include gravity waves, thermal tides, and planetary Rossby waves. Examples of therole of these processes in vertical coupling are presented in Sect. 4.2. The propagation con-ditions for all of these waves depend on the background winds, the thermal structure of theatmosphere, and characteristics of the wave such as wavelength and frequency. The linearwave equations, based on the assumption that the wave amplitudes are small, are often goodpredictors of wave propagation even when the conditions of linearity are violated.

Although most of the waves that are important in the atmospheric system are generated inthe troposphere, waves can also be forced in the stratosphere and mesosphere. Two examplesare the generation of the quasi-two-day wave by baroclinic instability (Plumb 1983; Pfister1985) and the forcing of the diurnal and semidiurnal tide by ozone heating (Hagan 1996).

2.2 The Stratosphere

2.2.1 Heating

The heating in the stratosphere occurs internally, unlike in the troposphere where much ofthe heating occurs at the interface with the surface. Most of the stratospheric heat comesfrom absorption of solar radiation by ozone. Although molecular oxygen is less importantthan ozone in the thermal budget, absorption of radiation in the Herzberg bands by O2 is thesource of O which in turn is the source of ozone (ozone is formed from the reaction of Owith O2). The solar ultraviolet radiation also photolyzes many other molecules and drivesthe active chemical system in the stratosphere.

Throughout the middle atmosphere, CO2 radiative transfer is the dominant cooling pro-cess. The increase in CO2 since the industrial revolution is changing the radiative balanceand is expected to lead to a continued cooling of the stratosphere. Water vapor and ozonealso make contributions to the cooling. Ozone changes affect both the heating and coolingrates. Changes in these three gases over recent decades have led to temperature trends in thestratosphere (Ramaswamy et al. 2001).

The stratospheric temperature structure is not in radiative balance. Heating occurs in thesunlit atmosphere whereas cooling occurs at all latitudes and all local times. There is nosolar heating in the darkness of polar night. However, heating due to dynamically drivenadiabatic descent occurs and keeps the polar temperature higher than that predicted fromthermal equilibrium.

2.2.2 Dynamics

The adiabatic heating that is so important in the middle atmosphere is the result of a wave-driven circulation. Waves propagate from the troposphere. According to theory, waves that

Interactions Between the Lower, Middle and Upper Atmosphere 5

are steady and not dissipating will not interact with the background atmosphere (Andrewsand McIntyre 1976). The waves that are most important for the stratosphere are planetaryRossby waves, gravity waves, and equatorially trapped waves such as Kelvin waves.

Planetary Rossby waves, which can normally propagate into the stratosphere only duringwinter (Charney and Drazin 1961), cause much of the variability in the wintertime middleatmosphere, including sudden stratospheric warmings (McIntyre 1982). Their propagationis affected by the background winds and temperature. They are forced in the troposphere andcan be dissipated by long-wave radiative transfer or by Rossby wave breaking (McIntyre andPalmer 1983). Quasi-stationary planetary Rossby waves are always present during the wintermonths. They are readily observed by satellite and are well-simulated in global climatemodels. Their amplitudes tend to be larger and more variable in the NH.

Gravity waves generated in the troposphere propagate into the stratosphere. Because theyare more difficult to observe and cannot be represented in most global climate models, theglobal impact of gravity waves on the stratosphere is less well-established. Dissipating grav-ity waves are responsible for the structure of the stratopause and polar night jet in the highlatitude winter (Hitchman et al. 1989; Siskind et al. 2007).

Although equatorially trapped waves have a limited geographical domain, their impactin the tropics is important. In the classic simulation of the quasi-biennial oscillation (QBO)of the zonal wind over the equator, Lindzen and Holton (1968) and Holton and Lindzen(1972) were able to generate the oscillation from interactions of equatorial Kelvin waves,mixed-Rossby-gravity waves, and other gravity waves with the background wind. More re-cent simulations by Kawatani et al. (2010a, 2010b) showed that the spectrum of waves thatcontributed momentum to the driving of a simulated QBO was quite broad. Waves fromplanetary scale, such as Kelvin waves, to short horizontal wavelength gravity waves all con-tributed to the generation of the QBO in the model.

The 2-dimensional mean circulation in the stratosphere is best viewed using the trans-formed Eulerian framework introduced by Andrews and McIntyre (1976). The transformedEulerian mean is a good approximation of the zonal mean Lagrangian mass flux. The trans-formed Eulerian mean meridional and vertical mass fluxes are determined by adding theStokes drift due to waves to the Eulerian mean meridional and vertical winds. For steadynon-dissipative waves, the Eulerian mean winds and the Stokes drift cancel and there isno net mass flux. A residual remains for dissipating waves. Thus a wave-driven circulationis one that owes its existence to dissipating waves. The mass circulation is known as thetransformed Eulerian mean or residual circulation. In the stratosphere, it is referred to as theBrewer-Dobson (B-D) circulation after pioneering inferences from observations of water byBrewer (1949) and of ozone column by Dobson (1956). In the lower stratosphere, the cir-culation goes from tropics to poles (Dunkerton 1978). Air enters the stratosphere throughthe tropopause in the tropics and slowly moves toward both poles. It can then reenter thetroposphere in middle or high latitudes of both hemispheres. In the upper stratosphere, theB-D circulation has a global cell; it moves air from the summer pole toward the winter pole.

One useful variable for assessing the B-D circulation is the age of air (Hall and Plumb1994). The age of air measures the time that has elapsed since an air parcel entered the strato-sphere. In the high latitude lower stratosphere the age can be several years. To determine theage of air from observations, studies have looked at the concentrations of long-lived anthro-pogenic gases that have a monotonic trend at the surface, such as CO2 and SF6. Althoughthese tracers work well for determining the age of air, they are not ideal for investigatingtrends that might have occurred as a result of climate change. Variations in the productionrate over the years make the analysis highly complicated and sometimes ambiguous (Garciaet al. 2011).

6 A.K. Smith

The tropopause is highest and coldest in the tropical region. Because cold air can holdlittle water, air entering the stratosphere is extremely dry. The seasonal cycle in the tropicaltropopause temperature (coldest during NH winter) gives a seasonal cycle in the humidityof the air entering the stratosphere. The vertical mixing is sufficiently low that patternsof higher and lower water mixing ratio can be seen ascending in the lower stratosphere.This phenomenon, called the “atmospheric tape recorder”, was first reported by Mote et al.(1996).

2.2.3 Composition

The distribution of the dominant molecules in the stratosphere is the same as in the tropo-sphere: molecular nitrogen and oxygen are the dominant gases and are well-mixed (Brasseurand Solomon 2005). The CO2 concentrations are similar to those in the troposphere but witha lag in the growth by up to 4–5 years. As mentioned in Sect. 2.2.2, the air entering the strato-sphere is very dry. Currently, roughly half of the water in the stratosphere enters through thetropopause and the other half is produced within the stratosphere from hydrogen releasedby the destruction of methane (CH4). The contribution of methane is believed to have beenlower in preindustrial times.

The most important trace gas in the stratosphere is ozone. Ozone protects the Earth’ssurface from ultraviolet radiation, provides most of the heating of the stratosphere by solarinsolation, and is highly active chemically. It has been extensively studied from satelliteusing infrared or microwave emissions, and ultraviolet absorption or scattering.

The production of ozone depends on the solar ultraviolet flux (which photolyzes molec-ular oxygen) and the temperature (which affects the kinetic rate of the reaction that makesozone from molecular oxygen and atomic oxygen). The production rate is therefore notstrongly influenced by global changes in the atmospheric composition. An exception to thisgeneralization is the impact of lower temperature due to greenhouse gas increases. Duringdaylit hours, ozone is lost by photolysis but is quickly regenerated when the O combines withO2 to form another ozone. However, there are a series of catalytic cycles that remove bothO3 and O, and hence cause a net loss of ozone. These involve the molecules and atoms inthe hydrogen, nitrogen, chlorine, and bromine families. Anthropogenic increases in chlorineand bromine are responsible for the severe springtime ozone loss in the lower stratosphereover Antarctica (the ozone hole), for ozone loss in the NH winter/spring lower stratosphere,and for ozone loss in the global upper stratosphere (WMO 2010).

Many of the reactions involved in the catalytic ozone-loss cycles have reaction rates thatdepend on temperature. The overall effect of a decreasing temperature, with no changes tothe background composition, would be to increase the ozone production rate and decreasethe ozone loss rate. The combination leads to a marked anti-correlation of perturbations intemperature and ozone. This applies not only for perturbations of temperature and ozonedue to waves on timescales of days to weeks but also for predictions of the response of theglobal ozone layer to greenhouse gas increases.

2.3 The Mesosphere

2.3.1 Heating

Like the stratosphere, the heating in the mesosphere occurs internally from solar ultravioletradiation. The dominant absorber is molecular oxygen. Another molecule that contributes iswater, which absorbs Lyman-α and Schumann-Runge radiation. The absorbed energy can

Interactions Between the Lower, Middle and Upper Atmosphere 7

photolyze the absorbing molecule, excite it to a higher state of energy, and/or contribute toincreasing its kinetic temperature. The energy is not necessarily available as heat at the timeof absorption.

Consider the absorption of ultraviolet radiation (<240 nm) by molecular oxygen. Thephotolysis gives two products: O(3P) and O(1D), i.e. atomic oxygen in the ground and ex-cited states. The photolysis uses much of the energy but the remainder is available as heat.The excited O(1D) atom can participate in one of many chemical reactions (O(1D) is short-lived and is much more reactive than O(3P)); it can collisionally transfer its energy to exciteanother atom or molecule; it can be collisionally quenched and convert the energy to heat;or it can spontaneously emit a photon and drop to the ground state.

The O(3P) atom has a lifetime that ranges from very short in the lower mesosphere tomonths in the upper mesosphere (Smith et al. 2010b). The reactions that either directly(reaction with O or O3) or indirectly (catalytic cycles with hydrogen-containing species)convert O(3P) to O2 are exothermic. This means, in simple terms, that the energy from thephoton that photolyzed O2 is not available as heat until the products recombine into anotherO2 molecule. Substantial transport can occur in the interim. For example, O generated in thelower thermosphere can be transported downward into the upper mesosphere and provide asource of heating even though the energy from the solar radiation did not penetrate to thispart of the atmosphere. The heating can be substantial during night due to the impact ofexothermic reactions. Some of the energy from O2 or O3 photolysis contributes to sponta-neous emission by O(1D) and other species (notably O2(1�) and vibrationally excited OH)as airglow. Much of the energy that is emitted as airglow is lost to space or to the ground anddoes not contribute to heating. The loss of energy through airglow is substantial (Mlynczakand Solomon 1993 estimate 20–40% for some processes) but is not precisely known.

The delay between the absorption of energy and the heating effect of the absorbed energyhas an additional impact on the dynamics. The net heating rate can shift away from havinga pronounced daytime maximum. The two exothermic reactions that have the largest impactnear the mesopause are O + O (i.e., the self-reaction of atomic oxygen) and O3 + H. Theformer is similar during day and night since O does not have a photochemically drivendiurnal cycle. The latter is actually larger during night because of the very large diurnalcycle in the concentration of O3 (Smith and Marsh 2005). One net impact of the exothermicreactions is to damp the magnitude of the diurnal cycle in heating in the upper mesosphere.

2.3.2 Dynamics

A defining characteristic of the mesosphere is the cold summer mesopause. The mesopauseat the summer pole is the coldest place in the Earth’s atmosphere. The mesopause is in-teresting for its cold temperature in summer and also because it is very far from radiativebalance (the summer is much colder than the winter). The cold summer temperature pro-vides an environment for polar mesospheric clouds (also called noctilucent clouds) eventhough the humidity is very low. The cold conditions are maintained by strong upwelling.The upwelling is part of a global circulation cell that moves air up at the summer pole,horizontally from the summer to the winter, and down at the winter pole (Smith et al. 2011).

The mesospheric circulation is driven by dissipating and/or breaking gravity waves(Lindzen 1981; Holton 1982). The waves are generated within the troposphere and thenfiltered by winds in the stratosphere. A gravity wave cannot propagate when the wave phasespeed is equal to the background wind. The point where the phase speed equals the back-ground wind is know as a critical layer. Because the stratospheric zonal winds are strongerthan the meridional winds, the critical layer filtering has the largest impact on waves propa-gating in the eastward and westward directions. The strong stratospheric jets in the solstice

8 A.K. Smith

periods filter out westward waves during the summer and eastward waves during the winter.The waves that are not filtered continue to propagate vertically but eventually break or dis-sipate in the upper mesosphere. The momentum deposition from the resulting asymmetricspectrum of gravity waves causes the stratospheric jets to reverse direction in the meso-sphere. The momentum deposition acts together with the Coriolis force to drive the summerto winter flow in both hemispheres.

The strong dependence of the mesospheric temperature distribution on waves from belowalso means it is susceptible to differences or changes in the regions below. An example ofthe impact of changes in the lower stratosphere is given in Sect. 4.2.

2.3.3 Composition

The composition in the lower mesosphere is similar to that in the stratosphere. One dif-ference is a shift in the catalytic cycles that destroy ozone. In the stratosphere, reactionsinvolving chlorine species and nitrogen species are very important in determining the ozonebalance. Beginning in the lower mesosphere, the catalytic cycles that destroy ozone arestrongly dominated by those involving hydrogen species. The reactive hydrogen species H,OH and HO2 are products of the destruction of methane and water.

Reactions involving species in the hydrogen and oxygen families continue to be impor-tant through the entire mesosphere. As noted above, many of the reactions are exothermicand they contribute to the distribution of heating. The proportion of heat from these reac-tions becomes larger in the upper mesosphere lower thermosphere (MLT) region. The com-position in the MLT differs from that in the stratosphere and lower mesosphere in severalimportant ways.

– The gases are no longer well-mixed but are beginning to separate based on their molecularmasses.

– Ions begin to make a contribution to the chemical system.– Some free radicals, for example hydroxyl (OH), become longer-lived and play a more

central role in the chemical system.– Atomic oxygen transitions from a very short lifetime below about 80 km to a long-lived

and abundant gas in the upper mesosphere.– Excited states of atoms and molecules affect the chemistry and energetics.– Ozone has a large diurnal cycle (smallest during daytime) due to the very rapid photolysis

rate.– Several important molecules, most notably CO2, are not in local thermodynamic equilib-

rium (LTE) in the mesosphere; see below.

Non-LTE conditions occur when low atmospheric density leads to a low rate of collisionsbetween molecules (Andrews et al. 1987). In general, strong cooling of the atmosphereoccurs through emission by vibrationally excited CO2. However, under non-LTE conditions,the low collision rate limits the exchange of energy from kinetic to vibrational levels. As aresult, the population of vibrational levels is not consistent with a Boltzmann distributionand the infrared CO2 cooling rates are substantially reduced over the LTE situation.

Several of the characteristics of the mesosphere (extremely cold temperatures, impor-tance of radical species and excited states) contribute to uncertainties in the chemical system.Laboratory measurement of the key reaction rates under realistic conditions is extremelydifficult. Satellite observations using atmospheric emissions are constrained by the verylow densities and therefore weak signals of natural emissions. Retrieval of temperature orconstituent densities using measured emissions from a non-LTE atmospheric state requires

Interactions Between the Lower, Middle and Upper Atmosphere 9

knowledge of the energy structure of the molecule. For all of these reasons, the detailedknowledge of mesospheric composition lags that of the stratosphere.

For example, observational estimates of one of the most important gases, atomic oxygen,differ substantially. As pointed out in Smith et al. (2010b), the O concentrations from differ-ent analyses are not compatible. O is derived during day and night using two different tech-niques and different radiance channels from the SABER instrument. When tidal variationsare taken into account, the concentrations are consistent between day and night, support-ing the reliability of the concentrations retrieved from SABER. However, O derived fromSABER is larger by about a factor of four than the concentrations derived from WINDII byRussell et al. (2004). The SABER values are also larger than the concentrations in the MSISempirical model (Picone et al. 2002).

3 Interfaces Between the Atmospheric Regions

3.1 The Tropopause

The tropopause has traditionally been identified as the altitude of coldest temperature or,equivalently, the altitude where the decline of temperature with altitude ceases. However,for certain purposes, it has been more useful to fine-tune the definition. The tropopausecan also be identified by isentropic surfaces or by potential vorticity (Holton et al. 1995).Chemically, the tropopause marks the location of sharp decreases in water and increases inozone. It is the point where vigorous vertical mixing ceases. The altitude of the tropopauseis higher in the tropics (equatorward of 30°) than at high latitudes. The altitude ranges fromabout 18 km near the equator to about 10 km at the poles and is discontinuous at ±30°.

In the tropics, rather than a distinct tropopause boundary, the transition region is known asthe tropical tropopause layer, or TTL (e.g., Fueglistaler et al. 2009). This takes into accountthe finite depth over which the transitions from troposphere to stratosphere occur. The TTLregion separates two regions which have quite different dynamics. Since the air that entersthe middle atmosphere passes through this region, its impact on transport of trace species isimportant.

The processes responsible for the tropopause were discussed by Held (1982). The trop-opause marks the transition between the vigorous dynamical mixing in the troposphere andthe radiative balance in the stratosphere. A general idea of the tropospheric thermal structureand tropopause location can be deduced using the concept of radiative-convective balance.Solar visible radiation passes through the atmosphere and is absorbed at the Earth’s sur-face (land or ocean). Sensible and radiant heat from the surface warm the atmosphere frombelow. A purely radiative balance is not in general possible in the troposphere because theheating from below, particularly over land surfaces, leads to convective instability near theEarth’s surface. The temperature profile of an atmosphere in radiative convective balance isconstrained to maintain stability, i.e., to have a lapse rate that is less than �d for dry air or�s for moist air. This concept has been used in simple radiative balance climate models formany decades. In these models, excess heat in the lower troposphere is redistributed in thevertical direction (as if by convection) until the profile achieves neutral stability.

As discussed in Sect. 2.2, the heating in the stratosphere comes from absorption of so-lar ultraviolet radiation by ozone. The maximum temperature at the stratopause occurs nearwhere the optical depth for the ozone heating (primarily the Hartley band) is unity. Thetropopause then marks the transition between the negative temperature gradient of the tropo-spheric radiative-convective system and the positive gradient of the radiative balance systemin the stratosphere.

10 A.K. Smith

Recent work has contributed to the understanding of the maintenance of the tropopause.The altitude and temperature of the tropopause are sensitive to the radiative balance in theupper troposphere and stratosphere and the stratospheric circulation (Thuburn and Craig2000; Shepherd 2002; Schneider 2004; Birner 2010).

Water in the middle atmosphere depends on the transport of both water and methaneacross the tropical tropopause. As air passes through the very cold temperature of thetropical tropopause, much of the water condenses into ice particles and settles out. This“freeze-dried” air has water concentrations of only a few ppm. The water concentration in-creases above the tropopause due to formation from byproducts of methane destruction. Themethane transport into the stratosphere depends on the tropospheric concentration and thenet mass flux across the tropopause; it is not directly affected by the tropopause temperature.

Observations indicate that the concentration of water in the stratosphere has varied overdecadal time scales. Part of the variation is a result of the increase of methane, which leadsto more water production by photochemistry within the stratosphere. However, the observedvariations in water are larger than can be explained by the methane increase and variationsin the tropical tropopause temperature. Water increased for several decades until about 2000and has decreased since then (Randel et al. 2006; Fujiwara et al. 2010) although the reasonsare still not understood. Stratospheric water is important because of its impact on the globalclimate. The infrared effects of stratospheric water increases during 1980–2000 can accountfor a acceleration of climate warming at the Earth’s surface (Solomon et al. 2010). Likewise,the water decrease since 2000 has slowed down the surface warming compared to what itwould otherwise be with unvarying stratospheric water.

Changes in the flux of water from the troposphere to the middle atmosphere also havean impact on the chemistry in the mesosphere (Marsh et al. 2003) and on the formation ofpolar mesospheric clouds (Thomas 1991).

3.1.1 The Stratopause

The transition between the stratosphere and mesosphere is defined by a change in the ver-tical gradient of the mean temperature from positive to negative. Although there are somedifferences between the stratosphere and mesosphere (see Sects. 2.2 and 2.3), the stratopauseregion is not normally a place of major transition in either dynamics or composition. This isthe reason that the terminology now often uses middle atmosphere rather than stratosphereand mesosphere.

The stratopause owes its altitude to two processes: the mean radiative balance and thethermal structure due to gravity waves. The radiative balance is due to ozone, which is pro-duced in the stratosphere. Ozone absorbs solar ultraviolet radiation in the Hartley, Hugginsand Chappuis bands. The largest heating occurs around 40–50 km; this is the mean altitudeof the stratopause. Away from the high latitude winter, the stratopause altitude is well pre-dicted by the radiative balance. This is approximately the altitude where the optical depth ofozone is unity. The maximum heating rate leads to a maximum in the temperature.

In the high latitude winter, there is no sunshine so the heating is small. In this region, thestratopause is strongly influenced by the adiabatic warming due to sinking motion (Hitch-man et al. 1989). This is particularly the case in the NH stratosphere, where the winterdynamics is more active. Gravity wave driving of the circulation, which becomes dominantin the mesosphere, can also affect the stratospause region. Gravity wave driving of the polarwinter circulation appears to be stronger in the SH.

Another quality that distinguishes the winter stratopause from the tropopause and meso-pause is the very large interannual variability in its location. Three well documented events

Interactions Between the Lower, Middle and Upper Atmosphere 11

during recent NH winters (2003–2004, 2005–2006, and 2008–2009) show a displacement ofthe stratopause to much higher altitudes. All three cases were characterized by the suddendisplacement upwards and slow return (over a period of weeks) for a region poleward of70°N (Manney et al. 2005, 2008, 2009). These events highlight the importance of dynamicsin controlling the stratospheric structure in winter high latitudes.

The poleward flow within the stratosphere is driven primarily by quasi-stationary Rossbywaves. The poleward flux is directly associated with downwelling motion in the polar winter.The stratospheric temperature is also affected by the larger-scale mean circulation that ischaracterized by downwelling through the entire winter polar middle atmosphere. Gravitywaves are a prime driver of this circulation in the upper levels (above 40–50 km). Thevertical propagation and eventual breakdown of gravity waves are strongly affected by thewinds through which they propagate.

3.1.2 The Mesopause

At the mesopause, the vertical gradient of temperature switches sign again, from negativeback to positive. The mesopause region is the coldest part of the Earth system. The overalltemperature is low because of a lack of heating. Molecular oxygen is the leading absorberof solar radiation.

In a climatological sense, there are two distinct mesopause positions. In the summer highlatitudes, the altitude is about 83–89 km (von Zahn et al. 1996) while at all other latitudesand times, its position is close to 100 km. Model simulations indicate that the lower andcolder position of the summer mesopause is due to very strong adiabatic cooling associatedwith strong upwelling.

On top of this mean description are several details. First, there is a difference in thesummer mesopause altitudes between the Northern (NH) and Southern Hemisphere (SH).The SH position is higher and not quite as cold (Xu et al. 2007). Polar mesospheric cloudabundances are lower in the SH (Bailey et al. 2007). Processes that have been proposed toaccount for this difference include the annual cycle in Earth-Sun distance (Chu 2003) anddifferences in gravity wave driving of the circulation due either to tropospheric sources orstratospheric filtering (for example, Dowdy et al. 2001).

Another feature of the mesopause is the existence of temperature profiles that containmore than one minimum. The “mesospheric inversion” or “double mesopause” has beendocumented in observations and models (Meriwether and Gardner 2000). Analysis by Sassiet al. (2001) and Salby et al. (2002) show that, in winter midlatitudes, the double mesopausephenomenon is associated with planetary-scale variations in dynamics. The existence ofplanetary waves in the stratosphere and mesosphere affects the filtering, propagation, andbreaking of gravity waves. The combined impact of planetary and gravity waves can givesharp changes in temperature and other dynamical fields in some phases of the planetarywave. From a ground-based observation, where only a single longitude can be observed,multiple “mesopauses” can appear in the vertical. Although single site observations of adouble mesopause are common, this appears to be solely a local phenomenon.

The mesopause is normally considered to be in the vicinity of the turbopause or ho-mopause (Chabrillat et al. 2002). The turbopause is the altitude where molecular dissipationof waves is more important than turbulent dissipation. Molecular diffusion processes sup-presse wave breaking and turbulence, hence the term turbopause. The definition of the ho-mopause is related: it is the altitude at which diffusive separation begins. Both the turbopauseand homopause represent a transition to conditions where collisions are not frequent enoughto keep the atmospheric constituents well mixed. Observations of wave properties (Offer-mann et al. 2007) and composition (such as CO2; see review by López Puertas et al. 2000)

12 A.K. Smith

suggest that there is a gradual change toward diffusive separation rather than an abrupt onsetfor either the turbopause or the homopause.

4 Coupling Between Regions

As discussed in Sects. 2 and 3, the basic thermal structure and the interfaces between at-mospheric regions are affected by dynamical and radiative processes through the depth ofthe atmosphere. In this sense, vertical coupling is the basis of the mean atmospheric struc-ture. However, here the term “coupling” is used to refer not to the maintenance of the basicstructure but to the response of the atmosphere to perturbations.

This section gives brief summaries of some examples of vertical coupling between atmo-spheric regions.

4.1 Radiative Coupling

The penetration of incoming solar radiation through the Earth’s atmosphere varies stronglywith wavelength. The absorption of the radiation depends on the concentrations of gases.Above the stratopause, the conversion of the absorbed energy to heat also depends on theconcentrations of gases that are involved in exothermic reactions or in quenching excitedstates. The interfaces between atmospheric regions (the tropopause, the stratopause, etc.) donot in themselves pose any barrier to the transmission of radiation.

The infrared radiation emitted by the Earth’s surface and by certain gases in the atmo-sphere can be reabsorbed and reemitted multiple times, particularly in the troposphere wherethe density of absorbing gases and aerosols is higher. Absorption can also be a heat sourcein the lower stratosphere. In the middle atmosphere, emission of infrared radiation by CO2

is the leading cooling process.Perturbations to the radiative balance can occur through

– changes in the solar flux (such as the 11-year solar cycle);– changes to the absorption of solar radiation (such as reduced absorption in the stratosphere

due to ozone loss);– changes in the reflection or absorption of radiation at the Earth’s surface (for example due

to changes in snow-cover);– changes in the reflection or absorption of radiation in the Earth’s atmosphere (for example

due to changes in aerosol or cloud amounts);– changes in the absorption of infrared radiation in the atmosphere (due to changes in water,

CO2, or other radiatively active gases);– changes in the composition in the upper mesosphere that affect the energy loss through

airglow.

The radiative changes can lead directly to changes in the thermal structure. Changes inthe thermal structure, particularly if the magnitude of the change varies with latitude, canlead directly to dynamical changes.

Examples are given in Sects. 4.1.1 and 4.1.2.

4.1.1 Impact of Antarctic Ozone Loss in the Stratosphere on the Troposphere andMesosphere

The primary heating in the stratosphere is absorption of solar ultraviolet radiation by ozone.Since 1980, the ozone over Antarctica during the spring has been substantially reduced from

Interactions Between the Lower, Middle and Upper Atmosphere 13

levels prior to that year. The cause of this ozone hole is well-understood; it is due to ozonedestruction by chlorine and bromine containing compounds that become particularly reac-tive during the special conditions of SH polar winter and spring. One effect of the ozoneloss is a reduction of absorbed solar radiation, which leads to reduced heating. The nega-tive trend in temperature in the lower SH stratosphere during winter is clearly evident inobservations (Thompson and Solomon 2002; Randel et al. 2009). Due to the thermal windrelation in the atmosphere (3), this polar temperature trend is associated with a trend in themean jet strength in the SH spring and early summer.

Much attention has been focused on the climate effects of the SH ozone loss (for example,Gillett and Thompson 2003; Baldwin et al. 2007; Son et al. 2010). A useful analysis tool forthis is comparisons of past climate simulations (hind-casts) and future climate predictionsbetween models with realistic ozone loss (and recovery in the future) and those without.These show that the effect of the ozone loss in the SH and of the future predicted ozonerecovery are felt all the way to the Earth’s surface. Models with realistic ozone loss (eitherthrough interactive stratospheric chemistry or through specification from observations) showsubstantially less warming over Antarctica during the 20th century than those that have non-varying ozone or that do not resolve the stratosphere.

Changes in the SH stratosphere due to the ozone hole can also affect the mesosphere.The primary mechanism is changes in the propagation of gravity waves. The ozone hole ledto trends in the lower stratosphere temperature and winds during late spring and early sum-mer. The trend in the zonal mean wind, as simulated in the Whole Atmosphere CommunityClimate Model, have affected the propagation of gravity waves into the mesosphere (Smithet al. 2010a). As a result, the model simulations show that the strength of the summertimemean circulation in the SH decreased at the onset of the ozone hole around 1980. The modelalso simulated warmer summer mesopause temperatures due to the change in gravity waveinteractions.

The previous example illustrates the interlinked nature of some of the coupling processes.In this case, a chemical change in the stratosphere (ozone loss due to increased chlorine andbromine) leads to a radiative change (less heating due to lower ozone). The radiative changeaffects the temperature which, through the thermal wind relation, affects the zonal wind. Thestratospheric wind change affects the propagation of gravity waves from the troposphere tothe mesosphere. The breaking or dissipation of these waves then affects the strength ofthe circulation in the mesosphere and, through changes in the adiabatic heating, the polarmesopause temperature (Smith et al. 2010a).

4.1.2 Impact of Stratospheric Aerosols on Tropospheric Climate

Volcanic eruptions can inject trace chemical species directly into the atmosphere. Oneof these gases is SO2. The sulfur that is injected into the troposphere washes out in amatter of days to weeks, so its direct climate impact has limited duration. Over a pe-riod of weeks, the sulfur injected into the stratosphere by explosive volcanoes becomesincorporated into sulfate (H2SO4) aerosol particles. The eruption of Mt. Pinatubo, thelargest in the satellite era, injected about 20 Mt of sulfur directly into the tropical lowerstratosphere. The circulation in the lower stratosphere (the Brewer-Dobson circulation dis-cussed in Sect. 2.2.2) tends to move air upwards in the tropics, poleward and then down-ward at the middle and high latitudes with a timescale of several years. Therefore, thesulfur injected by tropical volcanoes can persist for longer, become globally distributed,and can therefore have an impact over a wider area and longer period (Robock 2000;Stenchikov et al. 2006).

14 A.K. Smith

The best known climate impact of volcanic eruptions is cooling at the Earth’s surface,especially during summer. The cooling is caused by the radiative impact of aerosols. Theparticles absorb and reflect some of the solar radiation and therefore reduce the radiation atthe Earth’s surface. The net effect of eruptions on the surface temperature is difficult to deter-mine because of few eruptions and large natural variability but is estimated to be 0.1–0.2 K(Robock 2000). The stratosphere is a crucial piece in the climate response. The timescale ofseveral years for air to move through the stratosphere keeps the volcanic aerosols present inthe system.

The impact of volcanic aerosols on the heating also affects the stratospheric circulation.The enhanced heating in the lower tropical stratosphere can increase the B-D circulation(Tie et al. 1994). The altered circulation due to Mt. Pinatubo increased the strength of thetropical upwelling motion. This branch of the circulation brings up air that is low in ozone;therefore the net effect of a stronger upwelling was to decrease ozone in the tropical lowerstratosphere. There is also a photochemical response of ozone to volcanoes; the aerosolparticles enhance the rate of ozone loss due to chlorine and bromine.

The distribution of aerosols in the stratosphere varies irregularly over multi-year time-scales, due to intermittency in major volcanic eruptions and to variations in their latitude,time of year, and sulfur content. The impact that a given eruption can have depends on thetransport that distributes the volcanic products within the stratosphere. If the stratosphericcirculation accelerates due to climate change, as predicted by climate models (WMO 2010),this will affect the climate response to eruptions.

4.2 Coupling by Wave Processes

Vertical propagation of atmospheric waves can carry momentum and energy from one regionof the atmosphere to another. To maintain constant kinetic energy, wave amplitude growsexponentially with decreasing density. Under favorable conditions for propagation, a smallperturbation can grow to a large wave at a higher altitude. However, the converse impliesthat a perturbation at one altitude will have rapidly diminishing impact at lower altitudes.

As discussed above, waves will not interact with the background atmosphere unless theyare transient or dissipating (Andrews and McIntyre 1976). Sections 2.2.2 and 2.3.2 men-tioned the importance of planetary waves and Kelvin waves in the stratosphere and of gravitywaves throughout the middle atmosphere. While the waves themselves are generated in thetroposphere, the propagation and dissipation depend on the atmosphere through which theypropagate. The example of the impact of lower stratospheric changes on the wave processesin the mesosphere (Sect. 4.1.1) illustrates this. The message from this is that perturbationsdue to waves depend on the wave generation, the wave propagation, and the wave dissipa-tion.

Gravity waves cannot propagate beyond a critical line, where the background wind speedis equal to the phase speed of the waves. As discussed in Sect. 2.3, the critical layer effectleads to filtering of gravity waves by the mean stratospheric winds. This filtering then leadsto an asymmetric spectrum of gravity waves. When those waves break in the mesosphere,the momentum forcing drives the mesospheric winds and circulation (Holton 1982).

The winds in the winter stratosphere are never zonally symmetric because of the presenceof large amplitude planetary waves. An individual small scale gravity wave will propagatevertically through the local “background” wind, not the zonally averaged wind. This meansthat the filtering of gravity waves changes with longitude. When those waves break, themomentum forcing will also change with longitude. The variations in momentum forcingwill then lead to wavelike perturbations in the mesospheric winds. This mirror effect has

Interactions Between the Lower, Middle and Upper Atmosphere 15

been seen in observations (Smith 1996, 1997) and in numerical simulations (Smith 2003).In observations, it is manifest as an anti-correlation between waves in stratospheric windsand those in mesospheric winds.

Analysis has shown connections between variations in the middle atmosphere at northernand southern high latitudes. One of these phenomena is a relationship between the amplitudeof the semidiurnal tide in the summer hemisphere and the planetary waves in the stratosphereof the other hemisphere (Portnyagin et al. 1998; Smith et al. 2007; Murphy et al. 2009).Simulations by Yamashita et al. (2002) suggest that this is due to a propagation of the tidefrom the winter to the summer hemisphere. Although the tides are global modes, the detailsof their generation and vertical or horizontal propagation in an atmosphere with spatial andtemporal variations in temperature, winds, and heating are still not well understood.

Another connection between the hemispheres is a correlation between planetary waves inthe winter stratosphere and the polar mesospheric clouds at the summer mesopause (Karls-son et al. 2009). The processes that generate this link are not well understood. Analysisindicates that the signal probably goes from the stratosphere to the mesosphere of the winterhemisphere. Transfer of the signal to the opposite hemisphere may involve the mean circu-lation and/or impacts of wind changes on the propagation or breaking of the gravity wavesthat drive the circulation.

4.3 Coupling by Mass Exchange

Coupling among atmospheric regions can also occur by exchange of mass. One relevantprocess is the transport of chemically active gases from the mesosphere to the stratosphere.Another is the transport of water from the troposphere to the stratosphere.

As discussed in Sect. 3.1, the tropopause marks the top of the region of rapid verticalmixing. Movement of mass across the tropopause occurs much more slowly than verticalmovement of air within the troposphere. Flow across the tropopause is important for thecomposition of both the troposphere and the stratosphere. The impact on stratospheric waterwas discussed in Sect. 3.1. The advection of stratospheric air into the troposphere increasesthe concentrations of ozone and other gases (Holton 1990; Olsen et al. 2002).

Although the production of chlorofluorocarbons (CFCs) has been phased out by the Mon-treal Protocol, these molecules still persist in the atmosphere. The most abundant CFCs havevery long lifetimes in the troposphere but can be destroyed by photolysis and chemical re-actions in the stratosphere (WMO 2010). The slow removal reflects the slow rate at whichair from the troposphere is moved into the stratosphere and then returns to the troposphere.

Transport of mesospheric air into the stratosphere has received much attention in the lastdecade. A motivating factor for this is to assess the impact of reactive nitrogen (NOx, whichconsists of NO and NO2) on stratospheric ozone. Downward motion by the mean circula-tion during winter (see Sects. 2.2 and 2.3) can transport trace species with positive verticalgradients. Perturbations in the composition can occur if the timescale for transport is shortcompared with the times for mixing with the surrounding air or significant photochemicalproduction or loss.

Perturbations to NO (Randall et al. 2006, 2009), NO2 (Hauchecorne et al. 2007), CO(Jin et al. 2005; Funke et al. 2009), ozone (Smith et al. 2009), and OH (Damiani et al.2010) have been identified during periods of strong dynamical perturbation. It turns outthat the assessment of NOx is complicated because NO can also be produced in situ in thepolar mesosphere due to particle precipitation events (Randall et al. 2005). For this reason,NOx is not always a reliable indicator of enhanced transport between the mesosphere andstratosphere.

16 A.K. Smith

4.4 Observed Links with Unknown Mechanisms

For a number of decades, there have been studies reporting unexplained but statisticallysignificant links in the atmosphere. Many of these have dropped out of view as additionaldata contradicted the previous relationship. The most persistent have been associated withsolar variability, particularly the 11-year solar cycle.

The first-order atmospheric response to solar variability is fairly straightforward: an in-crease in the solar flux will increase the heating (assuming absorbing gases or surface prop-erties do not change) and will increase the photolysis rate of some compounds and cantherefore affect the composition. The direct heating can be estimated based on satellite mea-surements of the solar irradiance and knowledge of the absorption as a function of altitudein the atmosphere. The predicted changes in the solar flux penetrating the troposphere dueto the solar cycle observed or estimated over the previous century are quite small (IPCC2007). Likewise, the variability in the solar flux reaching the lower stratosphere is small solittle direct response is expected there.

Nevertheless, observational analyses show that various climate indicators are well-cor-related with the solar flux. Gray et al. (2010) give a comprehensive overview of the currentstatus of the observations and the search for understanding.

One correlation that has received much attention is a relationship between the 11-yearsolar cycle, the QBO in tropical winds, and the middle and high latitude winter strato-sphere. Analysis presented by Labitzke (1987) showed that the NH winter 30 hPa tem-perature was anticorrelated with the solar flux if one looked only at winters during whichthe QBO was in its westerly phase. Since then, many years of observations have been addedand the correlation still persists (Labitzke 2005). To date, there has been no explanationof the physical mechanism for the observed relationship. The explanation is confoundedby several factors. 1) The Holton-Tan relationship (Holton and Tan 1980) establishes alink between the tropical QBO and the high latitude stratospheric temperature. Separat-ing this mechanism from the solar cycle is not straightforward. 2) The two important vol-canic eruptions during this period, El Chichon in 1982 and Mt. Pinatubo in 1991, bothcame near solar maximum. These had an impact on the stratosphere that, with the shortduration of available observations, is not possible to separate statistically from the solar cy-cle. 3) The QBO is asynchronous with the annual cycle. The beating of these two cycles(the QBO at 24–30 months and the annual cycle at 12 months) gives a longer period vari-ation that is difficult to separate statistically from the solar cycle (Salby and Shea 1991;Smith and Matthes 2008). 4) Numerical models have not succeeded in reproducing the ob-served relationship.

The scientific community is divided about these correlation studies. Many researchersdismiss them as being spurious correlations. The thinking behind this is that, if one exam-ines enough time series (there are hundreds of possibilities) over what is, for an 11-yearcycle, a fairly short data record, there will be some that show correlations. The possibil-ity of spurious but significant correlations is well-established statistically. The atmospherehas several modes of longer term variability (ENSO, decadal oscillations, etc.) that in somecircumstances can be mistaken for a response to the solar cycle if extreme care is not taken.

Despite the reservations by many in the community, much effort continues to be ex-pended in looking for additional correlations and trying to explain the correlations that turnup.

Interactions Between the Lower, Middle and Upper Atmosphere 17

5 Conclusions

This document gives a brief description of atmospheric regions and some of the processesinvolved in their interactions. The coupling between regions has received much attention inrecent years as both observations have improved and numerical models have provided bettertools for analysis. The coupling described in this paper is categorized into three types ofprocesses: radiative, wave propagation, and mass exchange.

Vertical coupling occurs readily through radiative processes. Radiative processes canwork both ways; the less dense parts of the atmosphere can have an impact on the underlyinglayers. Several examples were presented for the effect of stratospheric changes (aerosoldistribution; ozone loss) on the troposphere. Note however that both of these involve localchanges to the stratosphere itself (such as its temperature) and therefore include dynamicalas well as radiative processes.

Vertical coupling through wave processes also occurs readily. Most of the waves that areimportant to the middle atmosphere are generated in the troposphere. The conditions forgeneration of the waves and their propagation through the intervening atmosphere both leadto remote effects.

The downward impact of waves was not discussed here. Because of the exponential de-crease in wave amplitude with increasing atmospheric density, downward propagation ofwaves is not expected to be significant. One proposal for downward wave coupling is theimpact of the middle atmosphere on the structure of planetary waves in the winter tropo-sphere. For example, recent studies (Perlwitz and Harnik 2003, 2004; Kodera et al. 2006)have shown that, under some circumstances, upward propagating planetary waves can bereflected in the stratosphere. This reflection has an impact not only on the wave structure inthe stratosphere but also on the wave within the troposphere.

Coupling of atmospheric regions through mass exchange is also possible. In this case,we focus on the concentrations of trace chemical species that are perturbed due to changesin the transport. The examples are water transport into the stratosphere and mesospheric airtransported into the NH polar winter stratosphere.

This overview gives a description of the regions of the atmosphere and a taste of someof the important coupling processes between them. This survey is not comprehensive butinstead reflects the experience and interests of the author. The knowledge of the atmosphereand its interactions is ever increasing due to increased measurements and improved numer-ical modeling capabilities. These will lead to better understanding of the interactions andbetter predictions of how they will affect the future of the atmosphere.

Acknowledgements The National Center for Atmospheric Research is sponsored by the National ScienceFoundation. Support for this work was also provided by the NASA Heliospheric Guest Investigator Program.

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