GEOPHYSICAL RESEARCH LETTERS, VOL. 25, NO. 20, PAGES 3819-3822, OCTOBER 15, 1998
Evolution of tropospheric ozone radiative forcing
D. S. Stevenson, C. E. Johnson, W. J. Collins, R. G. Derwent, K. P. Shine†
and J. M. Edwards
Climate Research, Meteorological Office, UK
Abstract. We present the first estimate of the evolution
of tropospheric ozone (O3 (T)) radiative forcing since 1860
and into the future. The UKMO 3-D chemistry-transport
model (STOCHEM) was used to simulate the tropospheric
composition in 1860, 1950, 1970, 1990 and 2100, by changing trace gas emissions. The future scenario used a doubled
CO2 climate. STOCHEM includes extensive non-methane
hydrocarbon (NMHC) chemistry, and produces a reasonable
simulation of present-day O3 (T). Radiative forcings caused
by the modelled changes in O3 (T) since 1860 were calculated
using the UKMO radiation code, and included clouds and
stratospheric temperature adjustment. Calculated changes
in the global annual mean forcing since 1860 were 0.13, 0.22,
0.29 and 0.48 W m−2 for the four years. Up to 1990 this
forcing scales linearly with the change in total NOx emissions since 1860; this linearity breaks down in 2100. The
1990 forcing is at the lower end of the range from previous
modelling studies (0.28 - 0.51 W m−2 ), but is still significant, enhancing the well-mixed greenhouse gas forcing by
over 10 %.
Introduction
Tropospheric ozone (O3 (T)) is an important greenhouse
gas, and is thought to have increased in concentration since
pre-industrial times (1860), particularly in the northern
hemisphere (NH) [Oltmans et al., 1998]. IPCC [1996] estimated that O3 (T) has contributed 0.2 - 0.6 W m−2 to the
radiative forcing since 1860, compared to 1.5 W m−2 from
CO2 . The wide range in this estimate reflects uncertainties
in the past and present O3 (T) distributions. These uncertainties arise partly because O3 is much more reactive than
the well-mixed greenhouse gases, and has a heterogeneous
distribution. The sparsity of reliable past measurements of
O3 [e.g. Marenco et al., 1994], preclude all methods of estimating past global O3 except modelling approaches. A
further difficulty is that O3 is not directly emitted, but is
formed as a secondary pollutant, during the oxidation of
CO and hydrocarbons in the presence of NOx. An additional major source of O3 (T) is transport from the stratosphere. In this study, an offline 3-D chemistry-transport
model (CTM) ran a suite of annual simulations, each using a different emission scenario, representative of individual
years between 1860 and 2100. Global O3 (T) fields were then
fed into a radiation code, generating radiative forcings due
to O3 (T) changes since 1860. All previous estimates of the
† Department
of Meteorology, University of Reading, UK
Copyright 1998 by the American Geophysical Union.
Paper number GRL-1998900037.
0094-8276/98/GRL-1998900037$05.00
O3 (T) forcing have concentrated on the total change since
1860; our study provides estimates of how the forcing has
evolved over this period. The time evolution is important
in, for example, model studies which attempt to simulate
the evolution of the climate response to the forcing.
Modelling
The offline chemistry and radiation models were driven
by archived meteorological data, generated by the UKMO
Unified Model (UM) at climate resolution [Johns et al.,
1997] (3.75◦ longitude x 2.5◦ latitude x 19 vertical levels).
Archived data comprise pressures, temperatures, humidities, winds, tropopause heights, and cloud, precipitation,
boundary layer and surface parameters. The CTM utilised
the data at 6-hourly resolution; the radiation code used
monthly-mean fields.
Chemistry-transport model
An updated version of the UKMO tropospheric CTM
(STOCHEM) [Collins et al., 1997] was used. This model
adopts a Lagrangian approach in which the atmosphere is
divided into 50,000 air parcels, advected using a 4th order
Runge-Kutta scheme with a 3-hour time-step. Within each
parcel, 70 species (including CH4 , CO, NOx , 10 hydrocarbons, and peroxyacetyl nitrate) compete in 174 photolytic,
gas phase, and heterogeneous chemical reactions, with a 5minute time-step. The chemical mechanism and photolysis
rates took part in the fast photochemistry model intercomparison [Olson et al., 1997]. The parcels’ positions and compositions are never re-initialised. Appropriately positioned
parcels receive emissions (Table 1) and are subject to dry
and wet deposition. At the top of the model (ca.100 hPa),
parcels receive an influx of stratospheric O3 (O3 (S)) and
HNO3 in areas of descent. Between advection steps, parcels
are mapped onto an Eulerian grid (5◦ x 5◦ x 9 vertical levels), used for inter-parcel mixing and data output. Vertical
mixing occurs in convectively active regions, by fully mixing
a fraction of the parcels in a column. The advection and
mixing schemes have been validated using 222 Rn tracer experiments [Stevenson et al., in press]. STOCHEM took part
in the Global Integration Modelling (GIM) intercomparison
[Kanakidou et al., 1998], producing a reasonable simulation
of present-day O3 (T) observations.
Five scenarios were run: 1860, 1950, 1970, 1990 and 2100,
each for 15 months, with the first 3 months discarded. The
only differences between the first 4 scenarios were the anthropogenic and biomass burning emission magnitudes (Table 2), and initial fields of CH4 , C2 H6 and CO from the
UKMO 2-D model [Johnson and Derwent, 1996]. The 2100
scenario used meteorological fields from a 2xCO2 version of
the UM [Senior, in press]. Each run used a present-day
O3 (S) climatology [Li and Shine, 1995]. In reality, O3 (S)
3819
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STEVENSON ET AL.: TROPOSPHERIC OZONE RADIATIVE FORCING
Table 1. 1990 emissions in Tg yr−1 , except NOx , which are
in Tg(N) yr−1 , and SO2 and DMS, which are in Tg(S) yr−1 .
See Collins et al. [1997] for details.
Biomass
Burning
Vegetation
Soil
Oceans
23.6
478.0
174.0
6.8
6.8
53.0
19.1
23.6
—
5.3
15.8
22.5
1.1
0.3
4.3
3.0
41.7
73.2
—
9.0
560.0
45.0
7.3
7.3
2.3
11.3
5.6
—
—
—
22.5
—
—
6.8
—
6.6
2.5
—
—
84.0
—
3.9
3.9
9.0
22.5
22.5
570.0
—
—
—
—
—
—
—
—
—
—
6.3
—
—
—
—
—
—
—
—
—
—
5.6
—
—
—
—
2.7
—
1.1
—
56.0
11.3
—
—
—
—
—
—
—
—
5.6
—
—
—
—
4.6
—
16.9
a Other NO emission sources are lightning and aircraft:
x
5.6 and 0.57 Tg(N) yr−1 .
b Other CH emission sources are animals, rice paddies,
4
tundra and wetlands: 96, 68, 56 and 73 Tg yr−1 .
c One other SO emission source is volcanoes: 11 Tg(S)
2
yr−1 .
and the spatial distribution of emissions vary through time,
with impacts on O3 (T). We assume these are second order
effects compared to changes in emissions’ magnitudes and
climate change impacts.
Altitude / km
NOxa
CO
CH4 b
C2 H6
C3 H8
C4 H10
C2 H4
C3 H6
C5 H8
O-Xylene
Toluene
H2
HCHO
CH3 CHO
Methanol
Acetone
NH3
SO2 c
DMS
Anthropogenic
16
12
8
4
0
-90
0
-60
4
-30
8
12
Latitude
16
20 24
ppbv
30
28
60
32
90
36
Figure 1. Zonal mean change in O3 (T), 1860 to 1990, for
July.
Convective cloud water and ice amounts were varied until
cloud radiative forcings were in reasonable agreement with
measurements [Harrison et al., 1990]. heating approximation [Shine et al., 1995].
An increase in O3 (T) reduces upward radiation from the
troposphere, cooling the stratosphere. Stratospheric relaxation occurs over a few months, whilst tropospheric adjustment takes decades, due to the sluggish oceanic response. A
better estimate of the long-term forcing on the troposphere
includes the stratospheric response [IPCC, 1996]. We iteratively adjusted stratospheric temperatures in the perturbed
cases, until stratospheric heating rates returned to their unperturbed (1860) values; this is the fixed dynamical heating
approximation [Shine et al., 1995].
Results and Discussion
Radiation model
The Edwards and Slingo [1996] radiation code was used
at low spectral resolution. The code produced results consistent with those from a recent intercomparison of responses
to O3 (T) [Shine et al., 1995]. In the stratosphere, O3 profiles were overprinted with present-day O3 (S) values [Li
and Shine, 1995]. Other gas profiles and aerosol burdens
were fixed at 1992 levels [IPCC, 1992]. Clouds were represented using distributions and dynamic cloud water and ice
amounts from the archive. Water droplets and ice crystals
were assumed to be spherical, with radii 7 µm and 30 µm.
Table 2. Emission factors (relative to 1990) for the five scenarios. These are applied uniformly globally. Sources: anthropogenic [Dignon and Hameed, 1989]; biomass burning based
upon population; aircraft [Reichow, 1990]. Also shown are the
global annual mean lifetimes (τ ) for O3 and CH4 .
Year
Anthropogenic
1860
0.0
1950
0.31
1970
0.8
1990
1.0
2100a 3.4,2.1,1.4
Biomass
Burning
Aircraft
NOx
τO3
/days
τCH4
/years
0.2
0.58
0.77
1.0
1.4,1.7,1.5
0.0
0.0
0.45
1.0
2.5
27.0
25.6
23.9
22.2
17.9
8.0
9.2
9.3
9.3
10.3
a The year 2100 scenario followed IS92a [IPCC, 1992]; the
factors reported are split into NOx , CO, and hydrocarbons.
This scenario also used a 2xCO2 meteorology.
Ozone fields for the five scenarios were produced. Figure 1 shows the zonal mean change in O3 (T) 1860-1990, for
July. Increases of over 30 ppbv occur in the upper troposphere, at mid-latitudes and polewards. Peak increases at
the surface are about 25 ppbv at 40-60◦ N, the main site
of O3 precursor emissions. Increases of over 5 ppbv are
seen almost throughout the troposphere. Changes are in
line with some observational data [Marenco et al., 1994],
and are similar to previous model estimates, e.g. [Berntsen
et al., 1997]. However, some measurement sites [Oltmans et
al., 1998] show decreases in O3 (T) over the past 3 decades;
these are missed by the model and may reflect our neglect
of changes in O3 (S) and the spatial distribution of emissions. Global annual mean FO3(T ) values since 1860 of 0.13,
0.22, 0.29 and 0.48 W m−2 were found for 1950, 1970, 1990
and 2100. Figure 2 shows the forcing for 1860-1990; similar spatial distributions were found for the other years, reflecting the fixed emissions distributions. The forcing has a
strong hemispheric and seasonal cycle. In the NH, O3 (T)
generally shows a spring/summer maxima; LW forcing is
maximised when the surface-to-tropopause temperature difference is greatest; hence FO3(T ) peaks in summer. Peak
forcings occur over N. Africa/Saudi Arabia, where surface
temperatures and albedos are relatively high, and the European O3 plume is present. During winter, forcing peaks
over the relatively warm oceans in NH mid-latitudes. Peak
SW forcing is seen over the high albedo Arctic in summer.
In the SH tropics, emissions from biomass burning dominate
3821
STEVENSON ET AL.: TROPOSPHERIC OZONE RADIATIVE FORCING
90N
45N
0
45S
90S
180
90W
0
90E
180
Mean = 0.29 W m-2
0.1
0.2
0.3
0.4
0.5
0.6
0.7
Figure 2. Annual mean FO3(T ) for 1990, relative to 1860.
O3 production and forcing. NOx from aircraft is responsible
for 6 % of the 1990 global FO3(T ) . These hemispheric signals combine to produce a peak global mean forcing in NH
summer, and a minimum in NH winter.
Evolution of FO3(T ) since 1860 (Figure 3) shows that 44 %
of the forcing occurred before 1950, and 78 % before 1970.
These fractions closely match the factors applied to emissions (Table 2), and indicate an almost linear response of
Figure 3. Evolution of global annual mean FO3(T ) (). After 1950, forcings are joined linearly (dotted line); prior to 1950
(dashed line) the evolution has been estimated by assuming linearity with the change in total NOx emissions, using data
from Dignon and Hameed [1989]. The bars show seasonal extremes for each year. Results from other models for 1990 are
also shown: (×) 2-D models [Hauglustaine et al., 1994; Forster et al., 1996]; (4) 3-D models [Lelieveld and van Dorland,
1995; van Dorland et al., 1997; Roelofs et al., 1997; Berntsen et al., 1997; Brasseur et al., 1998].
3822
STEVENSON ET AL.: TROPOSPHERIC OZONE RADIATIVE FORCING
FO3(T ) to emissions increases. This linearity has been used
to estimate the evolution of the forcing between 1860 and
1950, using estimates of NOx emissions from Dignon and
Hameed [1989] (Fig. 3). However, this relation breaks down
in 2100, suggesting that the troposphere’s capacity to generate O3 from NOx will gradually decline. This is partly
due to saturation of O3 production, and partly because a
warmer wetter troposphere will promote O3 (T) destruction,
reduce its lifetime (Table 2), and tend to mitigate the effect
of rising precursor emissions. Whilst the lifetime of O3 (T)
falls by 18 % between 1860 and 1990, the CH4 lifetime increases by 16 % (Table 2). This is because the increase in
radical concentrations outpaces the increase in O3 , but the
opposite is true for CH4 . These trends continue in the future
scenario.
These results fall at the lower end of the range in previous estimates of FO3(T ) (Fig. 3). One of the estimates
[Lelieveld and van Dorland, 1995] is for clear skies, and another [Hauglustaine et al., 1994] is an instantaneous forcing;
we find that neglect of clouds and stratospheric adjustment
both lead to an overestimate in FO3(T ) by ca.20 %; Berntsen
et al. [1997] reached a similar conclusion. Including these
reductions, the range amongst 3-D models is 0.28 - 0.42 W
m−2 . Probably the largest uncertainty is the change in midlatitude and tropical UT O3 .
Conclusions
Increases in anthropogenic and biomass burning emissions since 1860 have caused modelled O3 (T) concentrations
to increase (Fig. 1). We calculate that this extra O3 caused
a radiative forcing of 0.29 W m−2 in 1990, towards the lower
end of the range of estimates from other models. Nevertheless, FO3(T ) is still significant, enhancing the well-mixed
greenhouse gas forcing since 1860 by over 10 %. The evolution of the forcing (Fig. 3) has closely followed global NOx
emissions, although this linearity is expected to break down
in future, as O3 (T) production tends to become saturated
with respect to NOx , and the lifetime of O3 reduces as temperatures, humidities, and radical concentrations rise.
Acknowledgments. We are grateful to the UK Department of Environment, Transport, and the Regions, for its help
and encouragement through contracts EPG 1/3/93 (Air Quality
Division) and PECD 7/12/37 (Global Atmosphere Division).
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PO Box 243, Reading RG6 6BB.
(Received March 25, 1998; revised July 7, 1998;
accepted July 16, 1998.)