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The chemical products of aircraft jet fuel combustion are emitted at the engine nozzle exit plane as part of a high-velocity plume. This gaseous and particulate stream is subject to chemical and dynamical processes that influence downstream composition. Eventually, plume constituents irreversibly mix with, and are diluted by, ambient air. Subsequently, some of the emitted species act in concert with other natural and anthropogenic chemicals to change ozone abundances in the Earth's atmosphere. The ultimate fates of these aircraft-derived species are determined by larger-scale chemical and transport processes.
Concerns about NO and NO2 (i.e., NOx) emissions from present-generation subsonic and supersonic aircraft operating in the upper troposphere (UT) and lower stratosphere (LS) were raised more than 20 years ago by Hidalgo and Crutzen (1977) because these emissions could change ozone levels locally by several percent or so. Despite extensive research and evaluation during the intervening years, WMO-UNEP (1995) concluded that assessments of ozone changes related to aviation remained uncertain and depended critically on NOx chemistry and its representation in complex models. Because of large uncertainties in present knowledge of the tropospheric NOx budget, little confidence has been placed in previous assessments of quantitative model results of subsonic aircraft effects on atmospheric ozone. Assessment tools and their input data continue to improve, however, and reconsideration is appropriate in the light of the extensive research results published since the WMO-UNEP (1995) assessment.
The research results published since the WMO-UNEP (1995) assessment have addressed a number of issues relevant to the assessment of ozone impacts of present aviation. These issues have included the development of improved aircraft NOx emission inventories, updating of evaluated chemical kinetic and photochemical databases, studies of aircraft plume chemistry, and the development of three-dimensional (3-D) modeling tools. Reviews have also been published of U.S. (Friedl, 1997) and European (Schumann et al., 1997; Brasseur et al., 1998) research programs addressing ozone and other environmental impacts of present aviation.
In this chapter we evaluate, from a qualitative and quantitative standpoint, the impact on atmospheric ozone of aircraft exhaust species, emitted either directly from engines or produced as secondary products of processes occurring in aircraft plumes. Our evaluation is based primarily on global model calculations rather than ozone trends because expected changes are not easily discerned from observations, as discussed below. We use intermodel comparisons and atmospheric observations of ozone to test the physics and chemistry parameterized in these global models and identify areas of remaining uncertainty.
Most present-day jet aircraft cruise in an altitude range (9-13 km) that contains portions of the UT and LS. Because these two atmospheric regions are characterized by different dynamics and photochemistry, the placement of aircraft exhaust into these regions must be considered when evaluating the impact of exhaust species on atmospheric ozone. Determination of the partitioning of exhaust into the two atmospheric regions is complicated by the highly variable and latitudinally dependent character of the tropopause (i.e., the transition between the stratosphere and troposphere). Comparisons of aircraft cruise altitudes with mean tropopause heights has led to estimates for stratospheric release of 20-40% of total emissions (Hoinka et al., 1993; Baughcum, 1996; Schumann, 1997; Gettleman and Baughcum, 1999).
Carbon dioxide CO2) and water vapor (H2O) are easily the most abundant products of jet fuel combustion (emission indices for CO2 and H2O are 3.15 kg/kg fuel burned and 1.26 kg/kg fuel, respectively). However, both species have significant natural background levels in the UT and the LS (Schumann, 1994; WMO-UNEP, 1995). and neither current aircraft emission rates nor likely future subsonic emission rates will affect the ambient levels by more than a few percent. Future supersonic aviation, on the other hand (which would emit at higher altitudes), could perturb ambient H2O levels significantly at cruise altitudes. Regardless of the magnitude of the aircraft source, CO2 does not participate directly in ozone photochemistry because of its thermodynamic and photochemical stability. It may participate indirectly by affecting stratospheric cooling, which can in turn lead to changes in atmospheric thermal stratification, increased polar stratospheric cloud (PSC) formation, and reduced ozone concentrations.
Aircraft water contributions, although relatively small in the troposphere, lead to the atmospheric phenomenon of contrail formation. Depending on the precise composition of contrail particles-which is largely determined by the specific processes occurring in the aircraft plume and by the ambient atmosphere composition and temperature-the particles may act as surfaces for a variety of heterogeneous reactions (K�rcher et al., 1995; Louisnard et al., 1995; WMO-UNEP, 1995; Schumann et al., 1996; Danilin et al., 1997; K�rcher, 1997; Karol et al., 1997). The participation of contrails in atmospheric photochemistry is further addressed in Section 2.1.3.
NOx constitutes the next most abundant engine emission (emission indices range from 5 to 25 g of NO2 per kg of fuel burned) (Fahey et al., 1995; WMO-UNEP, 1995; Schulte and Schlager, 1996; Schulte et al., 1997). With respect to ozone photochemistry, NOx is the most important and most studied component; its aircraft emission rates are sufficient to affect background levels in the UT and LS. Moreover, its active role in ozone photochemistry in the UT and LS has been well recognized (WMO-UNEP, 1985, 1995). A great deal of the recent scientific literature has focused on aircraft NOx effects, and this chapter neccessarily reflects that focus. Aircraft carbon moNOxide (CO) emissions are of the same order of magnitude as NOx emissions (i.e., 1-2 g kg-1 for the Concorde aircraft and 1-10 g kg-1 for subsonic aircraft) (Baughcum et al., 1996). Like NOx, CO is a key participant in tropospheric ozone production. However, natural and non-aircraft anthropogenic sources of CO are substantially larger than analogous NOx sources, thereby reducing the role of aircraft CO emissions in ozone photochemistry to a level far below that of aircraft NOx emissions (WMO-UNEP, 1995).
Emissions of sulfur dioxide (SO2) and hydrocarbons from aircraft, at less than 1 g kg-1 fuel, are significantly less than the more prominent exhaust species discussed above (Spicer et al., 1994; Slemr et al., 1998). Their primary potential impacts are related to formation of sulfate and carbonaceous aerosols that may serve as sites for heterogeneous chemistry. This possibility is discussed in Section 2.1.3. Non-methane hydrocarbon (NMHC) emissions may also contribute to autocatalytic production of HOx, provided that the reactivity of the NHMCs is sufficiently large relative to that of CH4 to overcome their numerical inferiority. However, model studies have indicated that volatile organic emissions from aircraft have an insignificant impact on atmospheric ozone at cruise altitudes (Hayman and Markiewicz, 1996; Pleijel, 1998).
Although jet exhaust spends a relatively short time in the immediate vicinity behind the aircraft, a number of important processes occur during that time that influence exhaust gas and aerosol composition, hence the ozone-forming or ozone-depleting potential of the exhaust. The near-field evolution of jet aircraft exhaust wake can be divided into three distinct regimes-commonly termed jet, vortex, and plume dispersion. The time scales associated with these regimes are 0-10 s for the jet, 10-100 s for the vortex, and 100 s to tens of hours for plume dispersion-the latter time period defining the effective "lifetime" of the aircraft plume. The jet and vortex regimes are closely related; they are initiated at the exit plane of the engine nozzle and ended by atmospheric shear forces at distances of approximately 10-20 km behind the aircraft (Hoshizaki, 1975; Schumann, 1994).
Several fluid dynamic models are now available to study wake dynamics-namely, two-dimensional (2-D) jet mixing codes (Miake-Lye et al., 1993; Beier and Schreier, 1994; K�rcher, 1994; Garnier et al., 1996) and codes that capture the jet/vortex interaction and vortex break-up (Quackenbush et al., 1993; Lewellen and Lewellen, 1996; Schilling et al., 1996), some of them using vortex filament methods combined with large eddy simulations (LES) (Gerz and Ehret, 1997).
The small spatial and temporal scales of exhaust species distributions in near-field wakes hamper a robust comparison of model simulations with in situ observations of exhaust effluents. Nevertheless, the dynamic models have been successful in explaining the few observations of near-field tracer concentration, temperature, and humidity (Anderson et al., 1996; Garnier et al., 1996; Gerz and Ehret, 1997; Gerz and K�rcher, 1997). The data and calculations reveal a strong suppression of plume mixing and dispersion during the vortex regime. Vortex systems are composed of cylindrical core regions, not well mixed radially and entraining only small amounts of ambient air. As a result, vortex plume temperatures and associated H2O concentrations may be well defined from fluid dynamic simulations and known emission indices. Within the vortex, high concentrations of exhaust species interact with each other and with small amounts of ambient gases and particles over a range of temperatures that differ from those in the background atmosphere. It is likely that some of the chemical interactions occurring in the vortex regime will influence the eventual composition of aircraft-derived aerosol particles and gases.
The plume dispersion regime begins after disintegration of the wake vortex and extends to an area where the primary exhaust gas concentrations (i.e., NOx, H2O, CO, CO2) are of the same order of magnitude as the corresponding ambient background levels. Results from modeling and observational studies of aged plumes (Karol et al., 1997; Meijer et al., 1997; Schlager et al., 1997; Schumann et al., 1998) show that most plumes mix with the background atmosphere according to a simple dilution law that can be approximated with a Gaussian plume model that includes estimated and measured atmospheric shear and diffusion parameters (Konopka, 1995; Schumann et al., 1995; Durbeck and Gerz, 1996). The key observables for these models have been ice particles in visible contrails and measured CO2 that serve as tracers of the plume mixing process. All of the studies have indicated that during the 10-20 hrs of plume dispersion, the plume cross-section may grow to 50-100 km in width and 0.3-1.0 km in height, with a corresponding exhaust species dilution ratio (R-the ratio of the plume mass to fuel mass) up to 108 as a result of ambient air entrainment. From analysis of more than 70 aircraft plume crossings by research aircraft in the North Atlantic flight corridor, Schumann et al. (1998) proposed that R can be approximated by R=7000 (t/t0)0.8, where t0 = 1s for 0.006 < t < 104 s. The relative rate, (dR/dt)/R, of ambient air entrainment into the plume is on the order of 10-3 s-1 in the first minutes of plume dispersion but decreases to on the order of 10-4 s-1 over a 1-2 hr period (Durbeck and Gerz, 1996).
In the plume dispersion stage, aircraft-derived gas and particle concentrations are still highly elevated over background levels, but they interact with large volumes of ambient species under temperature and pressure conditions of the background atmosphere. The composition and reactive characteristics of aircraft-derived particles fully evolve in the vortex and plume dispersion regions as a result of aerosol-precursor photochemistry and particle condensation, coagulation, and agglomeration processes. These particle-forming processes are described in further detail in Chapter 3. In addition, chemistry process model calculations indicate that a significant fraction of emitted NOx is converted to other reactive nitrogen (NOy) species in the plume dispersion region during the daylight (Karol et al., 1997; Meijer et al., 1997; Petry et al., 1998). Observations of NOy in aircraft plume compositions are consistent with these results (Schlager et al., 1997).
Most interactions between ambient ozone and ozone-controlling gases and aircraft exhaust occur in the days and weeks following emission. Dispersion of exhaust on regional and global scales is dictated by the same large-scale atmospheric dynamic processes that control mixing of other natural and anthropogenic sources of gases and particles. During that time, aircraft-derived gases and particles participate in the natural chemical cycles that control ambient ozone levels. The following subsections provide an overview of ozone chemistry.
Approximately 80% of atmospheric ozone resides in the stratosphere, where it is produced via in situ photochemistry occurring predominantly in the tropical middle stratosphere, albeit with significant contributions from mid-latitudes. Stratospheric circulation patterns transport ozone from the tropical stratosphere poleward and then downward from the mid-stratosphere predominantly in the winter hemisphere. Stratospheric ozone is not only transported but also destroyed via photochemical reactions over the whole stratosphere. In addition, about 7-25% of the total ozone mass (WMO-UNEP, 1985; Wauben et al., 1998) is transported to the extratropical troposphere; this type of transport occurs most intensively in the winter and in the Northern Hemisphere. Different transport modes correspond to different time scales, ranging from days to years.
Ozone formation and destruction rates increase with height and change with latitude in the stratosphere. Consequently, ozone "lifetime" decreases with height from about a year in the LS to minutes in the upper stratosphere. At the uppermost altitudes, ozone lifetime is sufficiently short that its abundance is in local photochemical equilibrium (WMO-UNEP, 1985).
At lower altitudes, ozone is not in photochemical steady-state, and ozone transport by air motions of various scales becomes increasingly important. The primary mechanism for mean global stratospheric transport is referred to as the Brewer-Dobson circulation, with rising motion in the equatorial belt of the LS and air mass spreading to the poles in the middle and upper stratosphere, with more intensive transport into the winter hemisphere.
In summary, stratospheric ozone distributions are determined mainly by atmospheric motions in the nightime polar regions, by a mixture of transport and photochemistry in the lower and middle stratosphere, and by photochemistry in the upper stratosphere.
Sources of ozone in the troposphere are more numerous than in the stratosphere, as are the photochemical reactions participating in ozone production and loss. Although in situ photochemistry is the dominant source of tropospheric ozone, downward flux of stratospheric ozone represents a significant source, especially in the UT and in winter over high latitudes. Removal of tropospheric ozone occurs predominantly by photochemistry, with some contribution from surface deposition.
The lifetime of ozone in the troposphere varies with latitude and altitude; in general it is of the order of 1 month, a value that is smaller than the time scale for transport between the Northern and Southern Hemisphere troposphere, which is typically about 1 year (WMO-UNEP, 1985).
As mentioned in Section 2.1.1.1, emissions from present aviation are injected near the tropopause. Dynamic, chemical, and radiative coupling between the stratosphere and troposphere are among the important processes that must be understood if we are to provide an adequate description and prediction of the impact of aviation on atmosphere and climate. Of special significance is the exchange of chemical species between the stratosphere and the troposphere. In the stratosphere, large-scale transport takes place via the Brewer-Dobson circulation, induced by momentum deposition by planetary gravity waves. This circulation is responsible for the observed difference between the stratospheric temperature and its radiative equilibrium value. However, this exchange involves a wide spectrum of scales ranging from large-scale ascent and descent via synoptic scales toward transport by waves, convection, and turbulence (Brewer, 1949; Holton et al., 1995; McIntyre, 1995). For the impact of subsonic aviation, the focus is on the exchange between the troposphere and the lowermost stratosphere (Hoskins et al., 1985). This part of the stratosphere is strongly coupled with the troposphere and is separated from the LS by a region enclosed between the 380 K and 400 K potential temperature surfaces (Holton et al., 1995).
For the transport, it is useful to distinguish between different regions of the globe:
In the tropics, upward transport occurs mostly through deep convection, though
small-scale vertical mixing by gravity waves might also play an important
role. The tropics are the regions on Earth in which the largest net upward
transport into the stratosphere occurs and which therefore directly influence
the composition of the global middle stratosphere. Mixing between the tropical
and mid-latitude lower stratosphere is influenced by the subtropical barrier.
Between the tropical UT and subtropical LS, however, transport along isentropic
surfaces is important (Minschwaner et al., 1996; Plumb, 1996; Volk et al.,
1996).
In mid-latitudes, the exchange between the troposphere and lowermost
stratosphere flows in both directions, with a somewhat larger downward component
(Siegmund et al., 1996). Most of this transport is related to the occurrence
of extra-tropical cyclones and blocking anticyclones. In cyclones, polar stratospheric
air is drawn into the troposphere while subtropical tropospheric air is drawn
into the stratosphere. The intermediate process of tropopause folding is followed
by re-establishment of the tropopause. There is evidence that upward and lateral
mixing of tropospheric air into the stratosphere remains limited to the lowest
few kilometers of the mid-latitude LS (Dessler et al., 1995; Boering et al.,
1996; Hintsa et al., 1998).
Around the polar vortices, the exchange in the stratosphere occurs along isentropic
surfaces, from the polar vortex toward mid-latitudes, by filaments torn off
from the vortex boundary. Vertical transport in the vortex itself mostly takes
place in the form of large-scale descent caused by radiative cooling. Horizontal
transport on the equatorward flank of the polar night jet is sharply coupled
with vertical transport associated with diabatic descent caused by radiative
cooling of warm air within the descending branch of the baroclinic circulation
(Pierce et al., 1993, 1994).
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