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Three-dimensional climatological distribution of tropospheric OH:

update and evaluation.

C. M. Spivakovsky1 , J. A. Logan1 , S. A. Montzka2 , Y. J. Balkanski3 , M. Foreman-Fowler4 ,

D. B. A. Jones1 , L.W. Horowitz5 , A.C. Fusco1 , C. A. M. Brenninkmeijer6 ,

M.J. Prather7 , S.C. Wofsy1 and M.B. McElroy1 .

1. Harvard University, Cambridge, MA.

2. NOAA Climate Monitoring and Diagnostics Laboratory, Boulder, CO.

3. Laboratoire des Sciences du Climat, France.

4. University of Colorado, Boulder, CO.

5. National Center for Atmospheric Research, Boulder, CO.

6. Max Plank Institute for Chemistry, Mainz, Germany.

7. University of California, Irvine, CA.

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Abstract A global climatological distribution of tropospheric OH is computed using

distributions of O3 , H2O, NOx , CO, hydrocarbons and cloud optical depth, specified from

observations. Concentrations of OH are computed by forcing the system of kinetic equations to

the periodic solution, with a period of 24 hours. The global annual mean concentration of OH is

1.16×106 mol cm−3 (integrated with respect to mass of air). Mean hemispheric concentrations

of OH are within 1%. While the global mean OH increased by 33% compared to that from

Spivakovsky et al. [1990], mean loss frequencies of CH3CCl3 and CH4 increased by only 23%

because a lower fraction of the total OH resides in the lower troposphere in the present

distribution. The value 277K used for determining lifetimes of HCFCs by scaling rate constants

[Prather and Spivakovsky, 1990], is revised to 272K. The present distribution of OH is

consistent within a few percent with the present budgets of CH3CCl3 and HCFC-22. For

CH3CCl3 , it results in a lifetime of CH3CCl3 of 4.7 years, including stratospheric and ocean

sinks with atmospheric lifetimes of 43 and 78 years, respectively. For HCFC-22, the computed

lifetime is 11.4, allowing for the stratospheric sink of 229 years. Industrial sources of CH2Cl2

are sufficient for balancing its budget. Observed levels of CH2Cl2 (annual means) suggest that

no correction of hemispheric abundances of OH is necessary if the rate of interhemispheric

mixing in the model is increased to the upper limit consistent with observations of CFCs and

85Kr. If this rate is at its lower limit, an increase of OH in the northern hemisphere by 35%

combined with a decrease in OH in the southern tropics by 60% is suggested by observations of

CH2Cl2 . However, such large corrections are inconsistent with observations for 14CO in the

tropics and for the interhemispheric gradient of CH3CCl3 . The confluence of all available tests

does not suggest significant discrepancies except for a possible underestimation of OH in the

tropics in winter by 15-20%, and an overestimation in southern extratropics by ∼ 25%.

Observations of seasonal variations of CH3CCl3 , CH2Cl2 , 14CO and C2H6 offer no evidence

for higher levels of OH in the southern than in the northern extratropics. It is expected that in

the next few years, the interhemispheric gradient and annual cycle of CH3CCl3 will be

determined by its loss frequency, allowing for additional constraints for OH on scales smaller

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than global. The distribution of tropospheric OH, along with those for O3 , H2O, NOx , CO,

hydrocarbons, for important J-values, and for derived concentrations of selected species such as

HO2 , H2O2 , CH3O2 ,CH3OOH, CH2O, NO, and NO2 are available in the electronic form from

the authors.

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1. Introduction.

Ever since Levy [1971] presented a model of tropospheric chemistry with the hydroxyl rad-

ical as the key species, efforts continue to estimate accurately its concentration in the tropo-

sphere [e.g., McConnell et al., 1971; Weinstock and Niki, 1972; Wofsy 1976; Singh, 1977a, b;

Warneck, 1974; Lovelock, 1977; Crutzen and Fishman 1977; Fishman et al., 1979; Makide and

Rowland, 1981; Logan et al., 1981; Chameides and Tan, 1981; Volz et al., 1981; Crutzen and

Gidel, 1983; Prinn et al., 1983, 1987, 1992, 1995; Khalil and Rasmussen, 1984; Fraser et al.,

1986; Thompson and R.J. Cicerone, 1986a, b; Spivakovsky et al., 1990; Thompson et al., 1990;

Crutzen and Zimmermann, 1991; Brenninkmeijer et al., 1992; Mak et al., 1992, 1994; Krol et

al., 1998]. The abundance of OH determines lifetimes for CH4 , CO and a variety of industrial

pollutants, but the quest for accuracy has roots beyond the need to estimate the lifetimes of these

gases. The chemistry of OH comprises tightly coupled, mutually-compensating reactions which

in effect provide a buffer against changes in precursors and rate constants. Decades apart, Levy

[1971], Logan et al. [1981] and Spivakovsky et al. [1990] derived similar estimates for the

abundance of OH in the troposphere despite considerable evolution in the understanding of both

the chemical mechanism and the characterization of precursors. Errors of 15-25% in the global

mean concentration of OH may signify major misconceptions about the chemistry or the abun-

dance of precursors of OH in the troposphere. At the same time, testing global models for OH

has been associated with uncertainties of a similar or larger magnitude, intrinsic to deriving an

estimate indirectly, from budgets of species for which reaction with OH provides the dominant

sink and the sources are believed to be known: CH3CCl3 [e.g., Singh, 1977a,b; Lovelock, 1977;

Makide and Rowland, 1981; Logan et al., 1981; Chameides and Tan, 1981; Khalil and

Rasmussen, 1984, Fraser et al., 1986; Prinn et al., 1983, 1987, 1992, 1995] and 14CO [e.g.,

Weinstock and Niki, 1972; Volz et al., 1981; Brenninkmeijer et al., 1992; Mak et al., 1992;

1994]. (See Thomson [1992] for a review of studies of tropospheric OH through the early

1990s).

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Significant recent developments have affected both direct and indirect methods for estimat-

ing the abundance of OH in the troposphere. Most notably, the budget of CH3CCl3 was

modified by a 18% decrease in the calibration of CH3CCl3 [Prinn et al., 1995], a decrease of

about 12% in the recommended rate constant for reaction with OH [Talukdar et al., 1992] and

the discovery of an ocean sink for CH3CCl3 [Butler et al.,1991]. Observations of Brennink-

meijer et al. [1992, 1993] and Mak et al. [1992, 1994], together with earlier measurements by

Volz et al. [1981], provided for the first time a comprehensive description of the tropospheric

distribution of 14CO. Initial interpretation of these measurements indicated significantly higher

concentrations of OH than those predicted by models or inferred from the budget of CH3CCl3 at

the time [Mak et al., 1992]. In addition, higher concentrations of 14CO in the southern hemi-

sphere than in the north led to the suggestion that concentrations of OH are significantly lower in

northern than in southern midlatitudes [Brenninkmeijer et al., 1992], whereas most models

predicted slightly higher concentrations of OH in the north [e.g., Crutzen and Zimmermann,

1991; Spivakovsky et al., 1990]. Increasingly, however, it has been recognized [Spivakovsky

and Balkanski, 1994; Mak et al., 1994] that concentrations of 14CO are as sensitive to rates of

transport in the atmosphere as to the abundance of OH, and that the initial interpretations of

observations of 14CO must be revised.

Recent developments affecting the computation of tropospheric OH include the suggestion

by Michelsen et al. [1994] of a non-negligible quantum yield of O(1D) for wavelengths between

312 to 320 nm, confirmed by rate measurements [Talukdar et al., 1998]; there have been changes

in recommendations for other key rate constants [DeMore et al., 1994, 1997], in particular, a

decrease of ∼ 20% in the rate for reaction of OH with methane. New observational data affords

a better definition of precursors for OH such as O3 , NOx (defined as

NO2+NO+2N2O5+NO3+HNO2+HNO4), CO and H2O. Reactions with non-methane hydrocar-

bons (NMHC), omitted in earlier studies because of lack of observations, can now be included.

The distribution of cloud cover, highly uncertain in the past, can now be constrained by global

climatology from satellite observations, ISCCP [Rossow and Schiffer, 1991].

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Remarkable advances have been made in measuring concentrations of OH [e.g., Eisele et

al., 1997; Mount et al., 1997; Tanner et al., 1997, Mather et al., 1997; McKeen et al., 1997;

Wennberg et al., 1994, 1998; Brune et al., 1998]. Calculations of OH during the Tropospheric

OH Photochemistry Experiment, constrained by concurrent measurements of precursors, yielded

estimates in agreement with observed values within 30% for clean air conditions [McKeen et al.,

1997]. While these efforts lessen the uncertainties in the current chemical mechanism and lead

to its improvement, one has to rely on models to provide an integrated measure of the oxidative

capacity of the atmosphere over large regions because of the extreme variability of OH in time

and space.

The goal of this paper is to present an up-to-date global distribution for tropospheric OH

(essentially an update of that provided by Spivakovsky et al. [1990], referred to as S90 below),

and to evaluate the computed distribution using available observations of tracers. The 3-D dis-

tribution of tropospheric OH (averaged over 24 hours) was computed as a function of O3 , CO,

NOx , hydrocarbons, water vapor, temperature, cloud cover, and the density of the overhead

ozone column, by solving the system of kinetic equations using the full chemical mechanism

[DeMore et al., 1997; Talukdar et al., 1997a,b, 1998; Horowitz et al., 1998, Atkinson et al.,

1997; Paulson and Seinfeld, 1992], and by forcing the system to a periodic solution, with a

period of 24 hours. Unlike studies that compute the distribution of OH as a byproduct of a

fully-coupled simulation of O3 , NOx , CO and hydrocarbons [e.g., Muller and Brasseur, 1995;

Roelofs and Lelieveld, 1995; Wang et al., 1998a,b,c], which may suffer from imperfections in

various aspects of the model such as transport, deposition, and emissions, this study specifies

distributions of precursors based on observations, i.e., according to our best present knowledge.

As an exception, the distribution of isoprene had to be specified based on model results because

an appropriate climatology is not available. The distribution of OH, archived monthly with a

resolution of 10° longitude, 8° latitude and 9 pressure levels, can be obtained in electronic form

from the authors. This archive also includes concentrations of precursors for OH, computed dis-

tributions of selected short-lived species and J-values.

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Simulations of CH3CCl3 , CHF2Cl (HCFC-22), 14CO, C2H6 , C2Cl4 , and CH2Cl2 , dis-

cussed below in the context of the constraints these observations pose for the computed distribu-

tion of OH, were performed using the HARVARD/GISS/UCI Chemical Tracer Model (CTM)

[e.g., Prather et al., 1987; Jacob et al., 1987; S90; Jacob and Prather, 1990; Balkanski and Jacob,

1990; Balkanski et al., 1992, 1993; Chin et al., 1996a, b; Koch et al. 1996; Wang et al.,

1998a,b,c]. The CTM uses the wind fields, surface pressures, temperatures and convective mass

fluxes recorded every 4 hours for one typical year of the GISS GCM II simulation [Hansen et al.,

1983], with a resolution of 5° longitude and 4° latitude; height is resolved in 9 layers using σ-

coordinates, with 7-8 layers representing the troposphere. Each 8° by 10° grid box for the

present distribution of OH comprises four grid boxes of the CTM.

We begin by presenting computed concentrations of OH, including characterization of pre-

cursors adopted for the calculation (Section 2). The sensitivity of OH to uncertainties in the

specification of various precursors is discussed in Section 3. Section 4 provides a revision of the

value for temperature used for scaling tropospheric lifetimes of hydrochlorofluorocarbons

[Prather and Spivakovsky, 1990] needed because a smaller fraction of the abundance of OH

resides in the lower troposphere in the present distribution than in the distribution from S90.

Observations of CH3CCl3 and HCFC-22 are used to test the global annual mean concentration

of OH in Section 5. We show that the present distribution of OH is consistent with current budg-

ets of these gases. We continue to explore means to test computed distributions of tropospheric

OH on scales smaller than global [cf S90; Brenninkmeijer et al., 1992; Mak et al.,1992, 1994;

Spivakovsky and Balkanski, 1994; Goldstein et al. 1995a] using observations of CH3CCl3 ,

HCFC-22, C2H6 , C2Cl4 , and CH2Cl2 (Section 6). The utility and limitations of observations of

14CO for testing concentrations of OH are discussed in Section 7. In Section 8 we review the

unique opportunities for testing computed distributions of OH that may arise in the next few

years due to the phasing out of emissions of CH3CCl3 in compliance with the Montreal Protocol

[S90, Ravishankara and Albritton, 1995].

2. The 3-D distribution of tropospheric OH.

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We begin by describing distributions of precursors for OH adopted in this study.

Ozone. We replaced zonal mean ozone columns used in S90 from TOMS-Version 5 by a

2-D climatology derived from TOMS-Version 7 for 1978-1992. Zonal means for the new values

are 3% lower at midlatitudes than those used in S90; the consequent impact on global mean OH,

however, is less than 1%.

The global distribution of tropospheric ozone was specified using the climatology

developed by Logan [1998, submitted to JGR], incorporating ozonesonde data, tropospheric O3

columns ("tropospheric residual") [Fishman et al., 1990; Fishman and Brackett, 1997], and sur-

face observations. The main improvement in the specification of O3 (as compared to that

adopted in S90) is the resolution of longitudinal gradients in the tropics. The revised climatol-

ogy reflects a prominent feature of the tropospheric residual, high concentrations of O3 over a

large area in the southern tropics during the biomass burning period, including that over the

Atlantic. There is no indication in the observational data of a comparable increase in O3 in the

northern tropics, despite the fact that similar amounts of biomass are believed to be burnt in the

two regions [Hao et al., 1990]. The sparseness of observational data for O3 in the tropics does

not allow for confirmation of the residual data in the northern tropics (see Logan, [1998]). The

new climatology of O3 results in an increase of 3% in mean tropospheric OH, as compared to

S90.

Water vapor. The distribution of water vapor was specified using ECMWF monthly clima-

tological means from 1986 to 1989 archived at NCAR [Trenberth, 1992]. The ECMWF mois-

ture fields for that period have undergone extensive comparisons with observations. Liu and

Tang [1992] used radiosonde soundings from a world-wide network, including 52 tropical sta-

tions, to evaluate the surface and column-integrated specific humidity. In addition, they com-

pared the latter over oceans with 25 months of satellite observations from SSM/I and found good

agreement over most of the area. The mean and standard deviation of differences between the

ECMWF and radiosonde data sets were 0.04 g cm−2 and 0.36 g cm−2 , respectively (the range of

measurements was 0.5-7 g cm−2). Significantly, the mean and standard deviation of differences

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between satellite and radiosonde data were similar in magnitude: -0.02 g cm−2 and 0.37 g cm−2 .

Large discrepancies appear to be confined to relatively small areas off the west coast of con-

tinents, where the ECMWF values are higher than observational data by up to a factor of 2.

The column-integrated mean humidity reflects mainly the amount of water vapor in the

lower troposphere. Soden and Bretherton [1994] evaluated ECMWF relative humidity fields

averaged over 500 to 200 mb using satellite observations from GOES for 60° S - 60° N and 150°

W - 0° (the 6.7-µm channel spectral measurements attributing the highest weight to values from

400 to 250 mb). They report relative differences between the ECMWF values and those from

GOES of "roughly 23% to 45% in the regions of subtropical subsidence, 0% to 23% over the

northern and southern hemisphere midlatitudes, and -23% to -45% over areas of deep tropical

convection". As discussed in Section 3, concentrations of OH above 300 mb depend little on

specific humidity.

In comparison to this study, tropical humidity in S90 was too low below 800 mb by 10-25%

and too high at 700 mb by 20-30% over the oceans. The change in the global average concentra-

tion of OH due solely to the revisions adopted for the distribution of water vapor was small (∼

3% increase).

Nitrogen oxides. The distribution of NOx (Table 1) was based on an analysis of aircraft,

shipboard and surface data for NO and NOx [e.g., Torres and Thompson, 1993; Carroll and

Thompson, 1995; Emmons and Carroll, 1997; Bradshaw et al., 1998]. In deriving vertical

profiles, we used the analysis of Bradshaw et al. [1998] who gridded NO data from the

NASA/GTE and AASE aircraft campaigns, supplemented by measurements from other cam-

paigns [e.g., Drummond et al., 1988; Kondo et al., 1993; Rohrer et al., 1997]. Although these

data are not sufficient in spatial or temporal extent to define a climatology for NO, they provide

a series of "snapshots" that show some consistent patterns; for example, concentrations of NO in

the marine boundary layer are very low, a few ppt, those from 4 to 6 km tend to be in the range

10-40 ppt over both oceans and continents, while values for 10-12 km are generally in the range

10 to 150 ppt. Over the polluted continents, boundary layer concentrations are much higher,

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from 100 ppt to several ppb, and the profiles are "C" shaped [e.g. Drummond et al., 1988], while

in remote areas, such as the Amazon, concentrations are much lower, 10-40 ppt [Torres and

Buchan, 1988]. Boundary layer concentrations in remote regions affected by biomass burning

are elevated compared to those removed from such influence. Based on the features seen in the

observations, we allowed for different profiles for NO over land and ocean, south and north of

30°N, and in the southern tropics between 60°W and 45°E, with higher concentrations over the

latter region in austral winter-spring based on observations from the TRACE-A mission. Higher

values were adopted in the boundary layer for certain regions, as discussed below.

The profiles selected for NO are compared in Figure 1 to vertical profiles obtained in vari-

ous parts of the world. The standard "land" and "ocean" profiles are identical for 6 km and

higher altitudes. In addition to the data shown in the figure (from which the selected profiles

were largely derived), the ocean profile for 30°-90°N agrees well with measurements from Stra-

toz III from the east coast of North America and the west coast of Europe in June 1984 [Ehhalt

and Drummond, 1988; Drummond et al. 1988]; mean NO values for 20°-70°N were only 20%

larger than the standard profile in Figure 1a for 3-8 km, and were about a factor of 2 larger for

10-12 km. Values of NO were as low as 20 ppt at 4-6 km over New Mexico in summer [Ridley

et al., 1994], similar to those over the oceans, while values above were factors of 2-3 higher than

those in Figure 1a and b, because of the bias in sampling near convective storms. Measurements

from TROPOZ II in January 1991 [Rohrer et al., 1997] close to Europe were within a factor of 2

or less of the standard profile in Figure 1a, but downwind of North America were much higher,

100-300 ppt for 2-6 km.

NO concentrations in the vicinity of Hawaii, in the western Pacific, and in the southern

Pacific in September/October are generally similar to or smaller than the standard profile

adopted for 30°N - 35°S (Figure 1c and d), while profiles from the western Pacific in winter are

generally somewhat higher than the standard profile (Figure 1d). NO values reported for flights

from Japan to Indonesia at 4.5 km were about 20 ppt in continental air in winter and summer, but

only 7 ppt near the equator [Kondo et al., 1993]. The standard profile for 30°N-35°S is similar

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also to mean profiles for these latitudes derived from measurements made as part of PEM-

Tropics [Sandholm et al., 1999] (Figure 1e). Mean values for NO from TROPOZ II data are

20-60 ppt for 4-8 km between 30°N and 30°S over the Americas, higher than the standard

profiles. The TROPOZ II data have higher NO values than measured on other aircraft expedi-

tions in other seasons. The NO profile selected for land in the tropics is based on longitudinal

transects from ABLE IIA over much of the Amazon [Torres and Buchan, 1988], while the data

profiles in Figure 1g are for the area around Manaus only; values in the dry season for this

region were higher than those in the wet season. We chose to use the NO profile for 30°N to

35°S for southern mid-latitudes prior to the availability of data from PEM-Tropics and TROPOZ

II. The measurements from PEM-Tropics south of 35°S tend to show smaller concentrations of

NO than the standard profile, by about a factor of 2 below 5 km, but by as a factor of 2-6 at 8 km

and above (Figure 1f). However, the TROPOZ II data from the west coast of S. America tend

to be higher than the selected profile for southern mid-latitudes, by up to a factor of 3. These

differences may reflect seasonality in sources of NOx , but data are insufficient to address this

question. The NO profiles selected for regions affected by biomass burning are based on rather

limited data available from TRACE-A (Figure 1h for the Atlantic Ocean, and Figure 1i for

adjoining continents).

The values of NOx corresponding to day-time observations of NO were derived solving for

a periodic solution (with a period of 24 hours), using the full chemical mechanism; in these cal-

culations, climatologies of O3 , H2O, CO and hydrocarbons described in this Section were used

to define average concentrations for NOx in the specified regions. The profile for NOx obtained

in this manner for regions of Africa and South America affected by biomass burning was used

over sub-Saharan Africa during the biomass burning season.

Concentrations of NO (and NOx) are highly variable in surface air over the northern con-

tinents. The aircraft measurements in Figure 1b, those from New Mexico, and those from Stra-

toz III (1-2 km) show mean NO values of 20-70 ppt, while values are even lower over the Arctic.

The value selected for the standard profile in Figure 1b corresponds to ∼ 200 ppt NOx . Concen-

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trations of NOx reported for remote locations in the United States and Nova Scotia are ∼ 100-300

ppt, while median values at more polluted rural sites in the Eastern U.S. and Canada are 1-2 ppb

in summer, with somewhat higher values in winter; mean values are somewhat higher than

median values [Carroll and Thompson, 1995; Emmons and Carroll, 1997]. In the boundary

layer over industrial regions of Europe and North America we adopted an abundance for NOx of

1 ppb in summer and 2 ppb in winter.

Concentrations of NOx adopted for this study are higher than in S90 by factors 2-4 in the

tropics over regions affected by biomass burning and over industrial regions of northern midlati-

tudes; in addition, concentrations for NOx below 800 mb over oceans are higher in the present

work, by factors ranging from 1.5 to 2. The changes adopted for the distribution of NOx resulted

in an increase of about 7% in the tropospheric mean concentration of OH as compared to S90.

Carbon monoxide. We used the CTM as an interpolator to obtain a smooth distribution of

CO consistent with observations. For this purpose, the inventory of emissions presented by

Wang et al. [1998a] was adjusted arbitrarily to provide satisfactory agreement with observations

of CO from NOAA/CMDL and other surface sites [Novelli, 1992, 1994; Fraser, personal com-

munication], and where available, with observations of the CO column [] (using the distribution

of OH from S90). The agreement between observations and the distribution of CO specified in

the calculation of the present distribution of OH for surface and column data is illustrated in Fig-

ure 2. Concentrations of CO used in this study are about 20% lower in the southern hemisphere

and in the northern tropics as compared to S90 [cf Manning et al., 1997], resulting in an increase

by about 7% in the average value for OH.

Hydrocarbons. The methane field was assumed uniform in each semi-hemisphere, with

values from south to north of 1645, 1655, 1715, and 1770 ppm, respectively [Dlugokencky et al.,

1994].

Isoprene provides a major sink for OH near the surface over land in the tropics and at

midlatitudes in summer [e. g., Greenberg et al., 1985; Jacob and Wofsy, 1988]. Concentrations

of isoprene, however, are highly variable in space and time, and measurements are few. Conse-

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quently, for isoprene we had to depart from the principle of specifying distributions of precursors

from observations. Instead, we used the distribution of isoprene simulated using the fully-

coupled chemical tracer model of Horowitz et al., [1998], with emissions of isoprene reduced for

taiga, tropical forests and grassland by factor of 3, and increased for mixed deciduous forests by

factor of 3. In these modifications of emissions we were guided partly by new information (pers.

comm., Guenther), partly by comparisons with available observations for vertical profiles

[Rasmussen and Khalil, 1988; Jacob and Wofsy, 1988; Ayers and Gilett, 1988; Andronache et

al., 1994; Guenther et al., 1996; Helmig et al., 1998; TRACE A]. Over large regions however

concentrations of isoprene may be in error by a large factor. Nevertheless, once they reach ∼ 250

ppt, OH is depleted by more than a factor of 2, as shown in Figure 3, and therefore uncertainties

in concentrations of isoprene above that level do not contribute significantly to those in the total

tropospheric column of OH. At isoprene concentrations above 500 ppt, the sensitivity of OH is

greatly diminished because at these levels of isoprene, production of OH is dominated by photo-

lysis of products of isoprene oxidation, such as CH2O and organic peroxides. As a result, a

larger uncertainty is associated with the spatial extent of concentrations of isoprene in the range

100-500 ppt, than with the accuracy of the specified values above 500 ppt. The lifetime of

isoprene at 500 ppt is on the order of a few hours, and therefore little errors are expected to be

associated with the uncertainties in the horizontal transport of isoprene in the CTM. However,

transport within the boundary layer may be almost instantaneous, and therefore the vertical

extent of significant levels of isoprene is sensitive to errors in the height of the boundary layer

and frequency of convection. Model results with the 4° by 5° resolution on σ-surfaces were

adopted for the 8° by 10° resolution on standard pressure levels using linear interpolation in log

of isoprene because of the highly unlinear dependence of OH on isoprene. Our calculation of

OH neglects the effect of transport of products of isoprene oxidation, e.g., methacrolein, methyl-

vinylketone, by forcing concentrations of such species to a periodic steady state at a point.

The inclusion of isoprene decreased the global tropospheric mean concentration of OH by

4%, and increased lifetimes of CH3CCl3 and CH4 by 5%. Wang et al. [1998c] found that the

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global mean OH increased by 10% if emissions of isoprene were neglected, however their esti-

mate accounts for simultaneous changes in O3 , NOx , CO and CH4 , whereas in our case these

distributions are fixed.

Apart from isoprene and methane, we considered 13 hydrocarbons included in the chemical

mechanism: alkanes (C2H6 , C3H8, C4H10, and C5H12), alkenes (C2H4 and C3H6), aromatics

(C6H6, C7H8, and C8H10) and oxygenated species (methanol, ethanol, and acetone). Using

observations, we developed concentration profiles for 4 latitudes, 2 seasons, and for land and

ocean, and determined which species needed to be included in the calculation of the global OH

field with a series of sensitivity studies.

Measurements of hydrocarbons are most abundant for northern mid-latitudes. Values

selected for the continental boundary layer were based primarily on the data from surface sites

shown in Table 2, and were chosen to represent background conditions; values above the boun-

dary layer in summer were based on data obtained near North Bay, Canada. Median vertical

profiles were calculated for each hydrocarbon species for selected geographic regions, using the

GTE data merges provided by Bradshaw et al. [1996]. The marine profiles were based on data

from the Arctic, from near Goose Bay, Canada, and from the eastern Pacific north of 21° N

(PEM-West A); for PEM-West A a filter of C2H6 > 750 ppt was used to select mid-latitude air.

Aircraft data from the western Atlantic (0-4?km) [Penkett et el., 1993] were also considered.

The summer profiles for northern mid-latitudes are more reliable than those for other latitudes

and season. The concentrations were chosen to represent typical conditions away from pollution

sources. Higher concentrations of alkanes were found on STRATOZ II and III and on some

early cruises in the Atlantic [Rudolph et al, 1995]; these were used to provide an alternative set

of profiles used in sensitivity calculations.

Winter profiles were selected for northern mid-latitudes in the same manner as for summer,

but the aircraft data were limited to the western Atlantic and the eastern Pacific (PEM-West B),

and a few profiles measured in the vicinity of California [Singh et al., 1988]. We used a filter of

C2H6>1000 ppt to select mid-latitude values from PEM-West B (for latitudes > 21° N), and used

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these values above 3 km for land and for the entire profile for ocean; this resulted in profiles

similar to those reported by Singh et al. [1988] for the alkanes, which are long-lived in winter at

these latitudes. Concentrations of alkenes in the marine boundary layer, which are short lived

even in winter, were taken from Rudolph and Johnen [1990].

Vertical profiles in the tropics were derived largely from PEM-West A (10-22°N) and B

(7.5-21°N) in the eastern Pacific, TRACE-A in the south Atlantic, Brazil and Africa, three

cruises in the Atlantic [Singh et al., 1988; Rudolph and Johnen, 1990; Koppmann et al., 1992],

two in the Pacific [Singh et al., 1988; Donahue and Prinn, 1993], and continental measurements

in Africa and Brazil [e.g., Zimmerman et al., 1988; Rudolph et al., 1992a]. Profiles above the

boundary layer in the northern tropics were based almost exclusively on the PEM data, supple-

mented by limited measurements on TRACE-A transit flights; in the southern tropics, they were

based on TRACE-A data for October, and on PEM-B data near the equator for February. Con-

tinental data were lacking for the northern tropics, so the southern data were used to estimate

concentrations in the wet season and dry season respectively. The continental data are much

cruder estimates of typical values as the measurements are for rather short time periods, usually

a few weeks at most. Relatively few vertical profiles were measured over the continents during

TRACE-A, but they appear reasonably consistent with the surface data from the dry season

reviewed by Rudolph et al. [1992a].

The profiles in Table 2 have lower concentrations of alkanes in June-October than in

December- February (the biomass burning season) in the northern tropics; concentrations of alk-

enes are about the same in the two seasons, except in the continental boundary layer. The

profiles in Table 2 are relatively conservative. Higher concentrations of alkanes were found on

STRATOZ III in June compared to PEM-West A in October at similar latitudes, and on some of

the Atlantic cruises compared with data from the Pacific, at least in northern winter [e.g.,

Rudolph, 1988, 1995]. The aircraft data for the southern tropics (TRACE-A) were obtained

towards the end of the biomass burning season, in October, and show higher values than those

we estimated for December-February, which are based largely on surface data and on the equa-

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- 16 -

torial profiles from PEM-West B. The TRACE-A data from October show somewhat higher

concentrations of C2H6 than those found on STRATOZ III in June, but lower concentrations of

C3H8 and C4H10, suggesting that the TRACE-A values are not unreasonably high to use for

the period June-October.

Profiles derived for southern mid-latitudes were based on measurements from cruises to

∼ 30°S [Rudolph et al., 1990; Koppmann et al., 1992] and vertical profiles from TRACE-A that

sampled mid-latitude air, based on backward trajectory calculatrions. These data sets gave lower

concentrations of alkanes than surface data from 70°S reported by Rudolph et al. [1992b], but

were thought to be more representative of middle latitudes.

The concentration profiles for hydrocarbons are most reliable for alkanes, and the reliability

decreases in the order alkenes, aromatics, and oxygenated species, simply because of the quan-

tity of measurements. Similarly, the profiles are better known for June to October than for

December to March, because of the timing of the aircraft campaigns, and for northern mid-

latitudes compared with any other region, because of the availability of surface data, and of

several sets of aircraft data. While the vertical profiles selected for the tropics are less well

defined that those for northern mid-latitudes, there appears to be reasonable consistency among

the tropical measurements as shown in Table 3 for alkenes.

We found that, at present levels, reactions involving aromatics have an impact of less than

2% on computed concentrations of OH over most of the globe [cf Houweling et al., 1998].

Given that aromatics add significantly to the complexity of the chemical mechanism, we chose

to exclude them from our computation. We allowed for two species of alkenes (ethene and pro-

pene), four alkanes (ethane, propane, and two lumped species, C4,5 alkanes and C6-8 alkanes),

and three oxygenated species (acetone, ethanol and methanol). However, these NMHC appear

to be present in the atmosphere at low levels, with an effect on the global mean concentration of

OH of 7% (decrease) [cf Donahue and Prinn, 1993; Houweling et al., 1998]. This result should

not be interpreted to imply that NMHC play a minor role in the chemistry of the troposphere.

The effect of hydrocarbons on OH was considered in this study only for the given distributions

-- --

- 17 -

of O3 and NOx derived from observations, without taking into account the role of NMHC in

regulating these distributions [cf. Wang et al., 1998c]. Inclusion of these species, accompanied

by a decrease in CO over most of the area south of 30° N, led to a decrease of 10% in the frac-

tion of OH residing below 700 mb as compared to S90, because of the larger weight of

temperature-dependent loss processes (enhanced at higher temperatures characteristic of the

lower troposphere). In the tropical upper troposphere, above 200 mb, concentrations of OH

increased by 50-150% due to photolysis of acetone [Singh et al., 1995], with little effect on glo-

bal mean OH.

Temperature. Temperature fields were specified using ECMWF monthly climatological

means from 1986 to 1989 archived at NCAR [Trenberth, 1992]. The replacement of GISS GCM

fields had little effect on the global mean OH as well as on hemispheric and semi-hemispheric

means.

Clouds. The ISCCP cloud climatology provides the global distribution of the total optical

depth of clouds combined with the height of cloud tops [Rossow and Schiffer, 1991]. S90 used

the distribution of cloud optical depth from GISS GCM II [Hansen et al., 1983]. In the tropics

model reflectivities were lower than observed by almost a factor of 2. In addition, model clouds

(recorded as 5-day means) extended on average to lower altitudes and did not exhibit the

observed correlations between the height of cloud tops and cloud albedo evident in the Stage C2

ISCCP data [Rossow and Schiffer, 1991]. We replaced GISS CTM clouds with the ISCCP

clouds. Unfortunately, the vertical distribution of optical depth was not available. We chose to

distribute the ISCCP cloud optical depth uniformly with height between the cloud top and 900

mb; cloud layers were assumed to be bound by the standard pressure levels at which concentra-

tions of OH were computed. We determined the fraction of the solar flux reflected by each cloud

layer using the radiative code of Prather [1974] with Henyey-Greenstein phase function, for

g=0.87 [Sobolev, 1975], assuming a black surface as the lower boundary.

Little error is expected to arise from uncertainties in the vertical distribution of optical

depth for clouds with low tops, or for optically-thin clouds, as well as above and below the cloud

-- --

- 18 -

deck (see Figure 4). The largest error is associated with cases involving a high thin cloud resid-

ing above a low thick cloud and masking its actual vertical extent. For a severe case of mis-

placement, when a cloud is extended up to 300 mb instead of, for example, 700 mb, concentra-

tions of OH are underestimated everywhere in the region between the assumed cloud top and

900 mb, with the largest error occurring just above the actual cloud top. For a cloud with a day-

time-average reflectivity of 0.4, e.g., a cloud of optical depth of 5 above the equator at the

equinox, this error is ∼ 50% (Figure 4), and the error in the total column abundance is 25%. In

addition, the fraction of the total column abundance of OH residing below 700 mb is underes-

timated by 6%. As discussed below, on average this fraction for the present distribution is lower

by 21% than in S90, and about half of this difference is associated with the replacement of GISS

GCM II clouds by ISCCP clouds. Therefore in principle, such cases of erroneously extending

the cloud to a higher altitude may have contributed to the upward shift in OH in the new distri-

bution. However, the analysis of Stage 2 ISCCP data (monthly means of optical depth averaged

with respect to reflectivity) shows that only a relatively small fraction of the clouds with high

tops are optically thick. Of all the clouds extending above 600 mb in the tropics, only about 30%

have day-time-average reflectivities higher than 0.2 and only about 10% have reflectivities

higher than 0.3; in the extratropics in summer the fraction of thick high clouds is even smaller

(Figure 5). The current distribution of OH was computed using 7-year averages of cloud

reflectivities derived from the Stage C2 data.

The replacement of GISS CTM clouds led to a decrease in concentrations of OH below 800

mb by 5-10% and a comparable fractional increase above 700 mb (except at high latitudes). The

annual global mean concentration of OH increased by 2%.

The computed distribution of OH. Zonally-averaged concentrations of OH for four seasons are

shown in Figure 6; distributions for 700 mb are presented for January and July in Figure 7, while

global, hemispheric and semihemispheric averages are presented in Table 4 and zonal means in

Table 5. Consistent with previous studies, highest concentrations of OH arise in the tropics;

strong seasonal variations are predicted for midlatitudes reflecting variations in sunlight and

-- --

- 19 -

water vapor. The global tropospheric mean concentration of OH is 11.6.105 mol cm−3

(integrated over the year with respect to mass of air from the surface to 100 mb between 32°S

and 32°N and to 200 mb outside of this region), i.e., ∼ 33% higher than in S90 (see Table 4).

Changes in reaction rates, primarily in the O(1D) quantum yields (12%) and in reaction of OH

with CH4 [DeMore et al., 1997; Talukdar et al. 1998] are responsible for about a 19% increase in

OH. The remaining difference is a result of competing effects of increases caused by changes in

distributions of precursors, as described above (∼ 24%), offset on average by decreases of about

4% and 7% due to inclusion of reactions involving isoprene and other NMHC, respectively.

In contrast to S90, the present distribution of OH reflects the influence of biomass burning

in the southern tropics from June to October over Africa, South America, and the Atlantic, and

from November to March over sub-Saharan Africa; higher concentrations are predicted for these

regions. As in S90, mean hemispheric concentrations of OH differ little (see Table 4).

The vertical distribution of OH differs significantly from earlier results as reflected in the

shape of the global annual mean profile (Figure 8). In the present distribution, only 33% of tro-

pospheric OH (integrated with respect to mass of air) resides below 700 mb as compared to 42%

in S90. For gases with loss frequencies inhibited at lower temperatures, this upward shift in OH

mitigates the effect of the increase in the abundance of OH. While the global mean tropospheric

abundance of OH increased by ∼ 33% as compared to that in S90, global mean loss frequencies

of CH3CCl3 and CH4 increased only by ∼ 23% (evaluated using current rate constants for both

distributions). Two factors contributed in almost equal measure to the redistribution of OH

within the tropospheric column. Inclusion of NMHC, accompanied by a decrease in concentra-

tions of CO over most of the globe (as described above), gave a larger weight to loss processes

for OH that are enhanced at higher temperatures and are thus more efficient in the lower tropo-

sphere; most of this effect is due to NMHC other than isoprene. At the same time, replacement

of GISS GCM II clouds with the observed clouds resulted in larger production of OH in the tro-

pics in the mid-troposphere and smaller production in the lower troposphere because in the tro-

pics ISCCP clouds are on average twice as reflective as those from GISS GCM II (and extend to

-- --

- 20 -

higher altitudes). These effects were partially offset by an increase in water vapor below 800 mb

in the tropics.

Concentrations of OH above 200 mb increased by 50-150% due to photolysis of acetone

[Singh et al., 1995], with little effect on global mean OH. We emphasize however that com-

puted concentrations of OH in the upper troposphere may need significant revisions because of

uncertainties in distributions of acetone and NOx . In addition, the assumption of a periodic

steady state for H2O2 , CH2O and CH3OOH may not be appropriate for the upper troposphere

due to significant influx of these species from the boundary layer associated with deep convec-

tion [Prather and Jacob, 1997; Jaegle et al., 1997].

3. Sensitivity of OH to uncertainties in specification of precursors.

The tropical troposphere plays the major role in oxidation of CH4 , CO and other industrial

compounds [Logan et al., 1981]. In the distribution presented here, 75% of the tropospheric OH

and 80% of the loss of CH4 is between 32°S and 32°N, with 62% and 77% of these below 500

mb, respectively. Consequently, the chemistry of OH in the tropics below 500 mb is the main

focus of the discussion below. The sensitivity of OH to the specification of precursors is con-

sidered only for the current chemical mechanism and present levels of relevant species. The dis-

cussion in this section draws heavily on concepts and results from earlier studies [e.g., Logan et

al., 1981; Ehhalt et al., 1990; Zimmerman and Poppe, 1993]

Table 6 shows relative changes in the global mean concentration of OH due to a scaling

(within ±50%) of distributions of various precursors. The less-than-linear response apparent in

Table 6 can be understood best in the context of production, loss and internal transformations of

the family of species, HOx: OH+HO2+CH3O2+H+CH3O+O2CH2OH, as a subset of a larger

family, odd-H: HOx+2H2O2+2CH3OOH+HNO4+HNO2 .

Sensitivity to changes in O3 , H2O, and NOx . Most of HOx and odd−H below 500 mb is

produced in the form of OH by reaction

O(1D)+H2O→OH+OH. (1)

-- --

- 21 -

Dominant loss processes for OH, such as reactions with CO, CH4 , and NMHC, represent main

routes for production of HO2 (either directly or through intermediate steps involving H, CH3O2

and CH3O), the key species for the secondary production of OH through reactions

HO2+NO→OH+NO2 , (2)

and

HO2+O3→OH+2O2 (3)

(see Figure 9). Major loss processes for HO2 occur through reactions with CH3O2 as well as

with itself, and represent an efficient self-destruction of HOx (consuming two molecules of HOx

per reaction). An alliance of two factors: (1) the second-order loss of HOx and (2) the

significance of secondary production of OH, is at the root of the less-than-linear dependence of

concentrations of OH on the rate of its primary production, through reaction (1).

As a result of the second-order loss of HOx, the role of secondary production of OH is

higher at lower levels of HOx, hence it is higher on average at higher latitudes where there is

less sunlight and water vapor (Figure 10). In the lower troposphere over tropical oceans (except

during biomass burning), reactions (2) and (3) play comparable roles in secondary production of

OH, together amounting to less than a half of primary production. In polluted regions affected

by biomass burning or industrial emissions, reaction (2) is not only more important than reaction

(3) but it exceeds primary production of OH. The sensitivity of OH to changes in the rate of pri-

mary production is vastly different for these two regimes of tropospheric chemistry characterized

by the role of the secondary production and determined by the abundance of NOx . In the low-

NOx regime, a 50% change in concentration of O3 or H2O leads to a 25-35% change in OH

(Figure 11, a and b). In contrast, the sensitivity of OH to changes in the rate of (1) is diminished

in the high-NOx regime, e.g., during the biomass burning season over land below 700 mb (Fig-

ure 11, e and f), when the rate of recycling within HOx through reaction (2) is higher by more

than a factor of 4 than the primary production of OH through reaction (1), with the former pro-

viding ∼ 80% of the total. Although relative changes in concentrations of O3 and H2O affect the

rate of reaction (1) roughly in the same way, on average the sensitivity of OH to changes in H2O

-- --

- 22 -

is somewhat higher than to those in O3 (see Table 6). This difference is evident in the high-NOx

regime; it arises from the particular roles that H2O and O3 play in determining the rate of (2).

Higher concentrations of H2O lead to a higher loss frequency of HO2 through the H2O-catalytic

self-reaction and thus to lower concentrations of HO2 , whereas an increase in O3 results in lower

concentrations of NO since O3 provides the main loss process for NO. However, whether

increases in the rate of (1) are due to increases in concentrations of H2O or O3 , in both cases,

resulting increases in production of OH are mitigated by decreases in the rate of (2). In the

upper troposphere where reaction (1) contributes little to production of OH (Figure 10), concen-

trations of OH are lower at higher levels of O3 (Figure 11, a and e); in the region of transition

around 300 mb concentrations of OH are insensitive to changes in O3 .

On average, as shown in Figure 10, the primary production of OH through reaction (1) con-

tributes 40-60% to the total production of OH in the tropical troposphere below 500 mb, with the

secondary production, through reactions (2) and (3), providing up to all but ∼ 15% of the total.

The loss processes for HOx result in production of the major reservoir species for HOx, H2O2

and CH3OOH (Figure 9). The photolysis of H2O2 provides most of the remaining production of

OH in the lower troposphere (Figure 10) while conserving odd-H (photolysis of CH3OOH con-

tributes to production of HO2 , via reaction with CH3O). As in the case of HOx, odd-H is

removed mainly by processes destroying two molecules of odd-H per reaction, such as reactions

of H2O2 , CH3OOH and HO2 with OH and deposition of H2O2; these contribute to a decrease in

sensitivity of production of OH to the rate of reaction (1).

The sensitivity of OH to the rate of (1) is also impeded by a cascade production of species

destroying OH initiated by reactions of OH with CH4 , NMHC and CO. Reactions with such

species, including CH2O, CH3OOH and H2O2 , contribute more than 20% to the total loss of OH

in the tropics below 500 mb.

The results of Liu and Tang [1992] suggest that errors in specified concentrations of water

vapor in the lower troposphere may be within 25% over most of the globe. Logan [1998] gives

an estimate of about 30% for the overall uncertainty of her climatology of O3 . The error in the

-- --

- 23 -

global mean concentration of OH associated with errors in distributions of O3 and H2O is not

expected to exceed 25% (see Table 6) even if errors in concentrations of H2O and O3 conspire

everywhere to change concentrations of OH in the same direction. However, since there is no

reason for a systematic bias in specified distributions, or a correlation between errors in

specifications of H2O and O3 in the lower troposphere, an actual error in the global mean con-

centration of OH due to errors associated with distributions of these species is likely smaller. It

is clear that for higher accuracy of computed distributions of OH it is more important to minim-

ize the uncertainties in concentrations of H2O and O3 in the pristine tropical lower troposphere

than those in polluted regions. Unfortunately, it is in the tropical troposphere that our

knowledge of the ozone distribution is the weakest [Logan, 1998]. Larger errors in concentra-

tions of water vapor above 400 mb discussed above are not expected to affect significantly local

concentrations of OH (Figure 11, b and f) because reaction (1) contributes less than 20% to the

production rate of OH over much of that region (see Figure 10).

On average, a relative change in the NOx distribution (within ±50%) results in a response

in OH that is about 3-3.5 times smaller (Table 6). In the high-NOx-regime in the tropics below

800 mb, as over land in the biomass burning season, a 50% change in NOx leads to an ∼ 25%

change in OH in contrast to a 5% change in OH in pristine regions (Figure 11, g and c). As dis-

cussed in Section 2, the climatology of tropospheric NOx is not well defined and errors may be

larger than a factor of 2. However, if concentrations of NOx for the biomass burning regions

below 500 mb (Table 1b) are scaled by factors of 2 and 3, the global mean concentration of OH

increases only by 2% and 3% respectively. This lack of sensitivity arises from a relatively small

volume of the troposphere affected on average by biomass burning as well as from a lesser

dependence of OH on NOx once the NOx levels exceed ∼ 300 ppt [Logan et al., 1981] (despite

the dominant contribution of reaction (2) to production of OH). At these levels of NOx the rate

of (2) becomes less sensitive to NOx since reaction (2) becomes the major loss process for HO2 .

At still higher levels of NOx concentrations of OH decrease as reaction of OH with NO2 contri-

butes significantly to the removal of OH. Similarly, a decrease in OH with increasing concentra-

-- --

- 24 -

tions of NOx in the upper troposphere over polluted regions evident in Figure 11g is a result of

the increasing importance of such loss processes for OH as reactions with NO2 , HNO4 and NO,

of which the first two reactions in addition constitute the loss of odd-H.

In our view, the largest errors in computed mean regional levels of OH associated with

specification of NOx may arise from uncertainties between 800 and 500 mb, where particularly

high concentrations of OH are predicted, including for unpolluted regions (see Figure 6). A fac-

tor of 2 decrease of NOx at these altitudes results in ∼ 10% decrease in OH in the tropics (Figure

11c) and ∼ 15% at midlatitudes (not shown). To ascertain further uncertainties in the computed

distribution of OH due to those in the specified distribution of NOx , we replaced the distribution

of NOx in Table 1 by that simulated by Wang et al. [1998b]. The resulting global mean concen-

tration of OH (integrated with respect to mass of air) decreased by 6%, the hemispheric mean for

the northern hemisphere changed little, but the hemispheric mean for the southern hemisphere

decreased by 10% due to a decrease by ∼ 7% in the tropics and by 22% in the extratropics. Con-

centrations of NOx from Wang et al. [1998b] in the southern hemisphere are lower by a factor of

2 over the oceans in the tropics and by factor of 4-5 at midlatitudes. The available observations

do not allow to resolve these differences. A significant underestimation of concentrations of O3

above 700 mb in the southern hemisphere by Wang et al. [1998b] may be an indication of

insufficient levels of NOx . On the other hand, the analysis of seasonal variations of CH3CCl3 ,

CH2Cl2 and 14CO at southern midlatitudes (Sections 6.2, 7) suggests that concentrations of OH

presented here may be too high by 15-25% in that region. In the tropics however observations of

14CO give no indication of excessive levels of OH and call for an increase in OH in winter (Sec-

tion 7).

Sensitivity to changes in CO. Reaction with CO accounts for less than 40% of the total loss

of OH below 500 mb in the tropics, and less than 60% outside the tropics (Figure 11). Reaction

with CH4 contributes 15-20% to the loss of OH south of 30°N in the lower troposphere, with a

smaller effect north of 30°N. The remaining significant loss processes, accounting for 30-50%

of removal of OH below 500 mb, include reactions with the products of methane oxidation,

-- --

- 25 -

CH3OOH and CH2O, reactions with other hydrocarbons and their products, with H2 , and with

H2O2 . The importance of the loss processes unrelated to CO lessens the sensitivity of OH

towards changes in CO. A change in CO by ± 50% globally modifies the global annual mean

concentration of OH by -14% and +23%, respectively. (The apparent asymmetry with respect to

increase vs. decrease is characteristic of an inverse dependence). The sensitivity of OH to

changes in CO increases with height (Figure 10, d and h) because of the increasing role of CO in

removal of OH (Figure 11): the temperature-dependent rate-constants for reaction of OH with

CH4 and other hydrocarbons falls off with height much faster than the pressure-dependent rate

constant for reaction with CO.

Sensitivity to changes in concentrations of hydrocarbons. As discussed in Section 2, the

most important uncertainty associated with isoprene loss of OH is the vertical extent of

significant concentrations of isoprene (above 100 ppt), and in our estimate this uncertainty for

the global annual mean concentration of OH is on the order of +5%.

If other NMHC are not included, the global annual mean OH increases by 7%; if their con-

centrations are increased by a factor of 4, the mean OH decreases by less than 15%. As dis-

cussed above, inclusion of NMHC increases loss frequencies of OH more in the lower tropo-

sphere contributing to the redistribution of OH within the tropospheric column. An increase in

concentration of acetone by a factor of 4 globally increases concentrations of OH above 200 mb

in the tropics by ∼ 50%, with little effect on the global annual mean (a less than 1% increase).

Little uncertainty is associated with the distribution of methane, which is almost uniform and

well represented by existing observations. However, the sensitivity of the global mean concen-

tration of OH to changes in CH4 is of interest because the global burden of CH4 more than dou-

bled since preindustrial times and has been increasing until recently [e.g., Dlugokencky et al.,

1998]. A ± 50% change in concentration of CH4 throughout the globe results in -10% and +14%

change in OH (Table 6)

Sensitivity to changes in ozone column. Of all the factors determining concentrations of

OH, the greatest sensitivity is associated with changes in the overhead ozone column (see Table

-- --

- 26 -

6). The exponential dependence of absorption on the optical path, with a major fraction of

O(1D) production occurring for optical depth exceeding 1.5, leads to an amplification of relative

changes in the rate of reaction (1) as compared to those in the ozone column. Fortunately, the

climatology of the ozone column is well defined. The documented reduction of column ozone

over 1979-1994 [McPeters et al.1996] may have affected directly the global mean concentration

of OH over that period by 3-4% (neglecting the ensuing effect on CO and hydrocarbons), given

that no significant trend in column ozone was found in the tropics and reported decreases for

midlatitudes are 4-6% per decade.

4. Revision of the value for temperature appropriate for rescaling

mean tropospheric lifetimes.

Tropospheric lifetimes of gases destroyed by reaction with OH with temperature-dependent

rate constants, expressed as k=Aexp−B/T for temperature T, are often estimated by relating their

rate constants to that of a reference species, e.g., CH3CCl3 . Prather and Spivakovsky [1990]

showed, using the distribution of OH from S90, that this method is accurate to less than 7% over

a wide range of B (from 0 to 2500 K) if rate constants are evaluated at temperature T=277 K.

However, their result has to be revised for the present distribution of OH because a smaller frac-

tion of total tropospheric OH resides in the lower troposphere: only 33% of OH is below 700 mb

as compared to 42% in S90 (integrated with respect to mass of air). As shown in Figure 12,

which is similar to the key figure from their work, T=272K is a more appropriate choice for the

present distribution resulting in a less than 5% error over the whole range of B. Alternatively,

errors can be made even smaller if temperatures 271K and 277K are used for the lower and

higher end of the range, respectively. This estimate of errors, as well as that of Prather and

Spivakovsky [1990], was obtained assuming uniform mixing ratios of tracers across the globe.

Their work showed that errors are increased by a few percent for tracers with a 2:1 ratio of hem-

ispheric mixing ratios (still uniform within a hemisphere). We emphasize that this method is

appropriate for long-lived species and that additional significant errors can be incurred for

-- --

- 27 -

short-lived (with lifetimes of several months) because of large gradients characteristic of their

distributions. For example, simulations of C2Cl4 , CH2Cl2 and C2H6 , discussed in Section 6 as

constraints for OH, revealed that actual lifetimes of these gases are longer than those obtained by

scaling the lifetime of CH3CCl3 by 36%, 28% and 19%, respectively.

Given the mean tropospheric lifetime of CH3CCl3 (with respect to reaction with OH), one

can obtain an estimate of tropospheric mean OH (averaged with respect to mass of air) using the

same approach by seeking to estimate a lifetime of an imaginary species with the rate constant

Aexp−B/T at A=1 and B=0. It follows from Figure 12 that for the present distribution, the error

in the mean global OH is negligible if the mean loss frequency of CH3CCl3 (with respect to

reaction with OH in the troposphere) is evaluated at 270K (at 272K the error is ∼ 4%).

5. Constraints for the global mean concentration of OH

imposed by the budgets of CH3CCl3 and HCFC-22.

In this Section, as well as in Sections 6 and 7, we will refer to Figure 13 summarizing

results of evaluation of the computed distribution of OH using observations of various tracers.

It shows ratios of the computed mean OH to the most likely value implied by observations of

various tracers, i.e., assuming the most likely values for the magnitude of sources, absolute cali-

bration, rate constant for reaction with OH and strength of other sinks; observational constraints

themselves are also taken at their most likely values (e.g., means over relevant time and space

intervals). Two kinds of uncertainties are specified. One is defined assuming that the mean OH

implied by observations of tracers is known exactly (solid lines); for example, uncertainties in

sampling concentrations of CH3CCl3 in the proximity of source regions in the model, or relevant

uncertainties in transport rates in the model would belong to this category. The other is shown

with respect to the ratio equal unity and is associated with determining the mean OH implied by

observations, i.e., with the magnitude of sources, absolute calibration or strength of other sinks

(dashed lines). Uncertainties associated with the variability in observations, while reflected in

figures for individual constraints, are not included in this summary. Ranges are computed for

-- --

- 28 -

the worst case scenarion, i.e., assuming that errors will conspire to affect the ratio in the same

direction.

Results of global simulations of CH3CCl3 are compared with observations [Prinn et al.,

1995] in Figure 14. The history of emissions of CH3CCl3 and their spatial distribution is from

Midgley and McCulloch [1995]. The stratospheric loss frequencies (Table 7) were obtained in a

manner described by Prather et al. [1987] and S90, using the 2-D model of Schneider et al.

[1998]. Simulations were initialized using ALE/GAGE observations [Prinn et al., 1995] for

January 1979. Solid lines represent the simulation with the computed distribution of OH

(referred to as standard OH below) and no ocean sink. The lifetime of CH3CCl3 in this simula-

tion is 5.0 years. Dotted lines correspond to simulations with OH reduced and increased by

25%. The rate of growth of calculated concentrations of CH3CCl3 using standard OH is some-

what higher than observed leaving room for an ocean sink of about 6% (the lifetime of CH3CCl3

in the simulation resulting in the observed long-term trend in CH3CCl3 is 4.7 years [cf Prinn et

al., 1995]).

Assuming estimates for uncertainties in emissions from Midgley [1989, 1992] and Midgley

and McCulloch [1995] (less than ±5%), and those for the absolute calibration from Prinn et al.

[1995] (less than ±5%), we conclude that the estimate of the mean lifetime of CH3CCl3 (4.7

years), with respect to all loss processes, is accurate to better than ±10%. This uncertainty may

be significantly reduced in the next several years, with emissions of CH3CCl3 being phased out

in compliance with the Montreal Protocol [Ravishankara and Albritton, 1995] as discussed in

Section 8.

The second tier of uncertainties in estimating the average abundance of OH in the tropo-

sphere using the long-term trend in CH3CCl3 arises from determining the role of loss processes

other than reaction with OH. Butler et al. [1991] give 59-128 years as a likely range for atmos-

pheric lifetimes of CH3CCl3 due solely to the ocean sink. The present distribution of OH allows

for the ocean sink of ∼ 78 years, i.e., near the middle of the range given by Butler et al. [1991].

The atmospheric lifetime of CH3CCl3 due solely to stratospheric loss in the model is 43

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- 29 -

years (computed by relating the loss of CH3CCl3 in the stratosphere to the total mass). We adopt

here an estimate for atmospheric lifetime of CH3CCl3 due to stratospheric sink of 34±7 y [Volk

et al., 1997] (somewhat shorter than the IPCC estimate of 45 years) and conclude that the global

loss of CH3CCl3 in the model is underestimated by ∼ 3%. Since errors in the stratospheric sink

in the model appear to result in underestimation of the global loss of CH3CCl3 by 3%, and the

simulation with the standard OH leaves room for the middle-of-the-range magnitude of the

ocean sink, we conclude that the mean loss frequency of CH3CCl3 for computed OH is about

3% higher than the most likely value for the mean OH implied by the observed long-term trend

in CH3CCl3 (Figure 13). Most of the stratospheric loss, due to photolysis and reaction with OH,

occurs above 70 mb. In the model, about 11% of global loss of CH3CCl3 occurs in the strato-

sphere where ∼ 8% of the total mass of CH3CCl3 resides; mean loss frequency in the model stra-

tosphere is ∼ 37% higher than in the troposphere. As discussed in Section 7 however, the tropo-

pause in the model is too high; as a result, regions of the stratosphere below 70 mb in the tropics

and below 150 mb in the extratropics, containing 25-35% of mass but incurring little loss of

CH3CCl3 , are excluded from the consideration of the mean stratospheric loss. In reality, there-

fore, the mean stratospheric loss is comparable to the mean tropospheric loss of CH3CCl3 due to

reaction with OH.

The third source of uncertainty in the estimate of the global mean tropospheric OH arises

from that in the rate constant for reaction of CH3CCl3 estimated at ±11% [Talukdar et al., 1992].

As shown in Figure 13, the combined uncertainty in the estimate of the tropospheric mean OH

(weighted by its loss frequency in reaction with CH3CCl3) is in the range from -24% to +27%.

Errors in excess of ±15% however are improbable, since potential corrections (in absolute cali-

bration, emissions, the rate constant, and non-OH sinks) are not expected to affect results in the

same direction. Since errors in the stratospheric sink in the model appear to result in underesti-

mation of the global loss of CH3CCl3 by 3%, and the simulation with the standard OH leaves

room for the middle-of-the-range magnitude of the ocean sink, we conclude that the mean loss

frequency of CH3CCl3 for computed OH is about 3% higher than the most likely value for the

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- 30 -

mean OH implied by the observed long-term trend in CH3CCl3 .

Figure 15 compares the observed and simulated rate of growth of concentrations of HCFC-

22 in 1992-1996 [Montzka et al., 1993, 1996; Miller et al., 1998]. The history of emissions is

from Midgley and McCulloch [1997] for 1992-1994, and AFEAS [1998] for 1995-1996. Simu-

lations were initialized from observations for January 1992. Stratospheric loss frequencies

(Table 7) were computed in a manner similar to that for CH3CCl3 , by integrating loss and mass

of HCFC-22 from 150 mb to 70 mb and from 70 mb to the top of the atmosphere using the

model of Schneider et al. [1998] with a vertical resolution of 2 km. Most of the stratospheric

loss of HCFC-22, mainly through reaction with OH, occurs above 10 mb, outside of the CTM

domain. The average loss frequency for 70-0 mb was applied to the 70-10 mb layer of the

model. The atmospheric lifetime of HCFC-22 in the simulation with standard OH (solid lines) is

11.4 years [cf Kanakidou et al., 1995; Miller et al., 1998]. The stratospheric loss gives an atmos-

pheric lifetime of 229 years. Dotted lines represent simulations with OH modified by ± 25%.

The standard simulation underestimates the rate of increase in HCFC-22 by 2-5%. The

unreported emissions, believed not to exceed 10% [P. Midgley, pers. comm., 1998]), were not

included. Uncertainties in calibration of measurements are estimated not to exceed ±5%

[Montzka et al., 1993], and those in the rate constant ±15% [DeMore et al., 1997].

Although significant uncertainties are associated with determining "true" global mean OH

using observed long-term trends of CH3CCl3 and HCFC-22 (dashed lines in Figure 13), the con-

sistent results for two independent constraints lends additional confidence to these estimates.

6. Constraints on the hemispheric and semi-hemispheric scale.

6.1. Annual mean levels of C2Cl4 and CH2Cl2 as a constraint for regional concentrations of

OH. Estimates of mean concentrations of OH on scales smaller than the global are required for

interpreting measurements and understanding sources of such species as CO, with the lifetime on

the order of a month in the tropics and in summer at temperate latitudes. There is a distinct

interhemispheric asymmetry in distributions of precursors for tropospheric OH. Species

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- 31 -

involved in both production and loss of OH are present at significantly higher concentrations at

northern than at southern midlatitudes [e.g., Logan et al., 1981]. The disparity extends to lower

latitudes. Concentrations of CO, the species providing the major sink for OH, are in general

higher in the northern than in the southern tropics because of the proximity of industrial regions

of Europe and North America. However, concentrations of O3 , the major source-species for

OH, appear to be lower in the northern tropics than in the south [Fishman et al., 1990] as dis-

cussed above. The larger land area in the northern hemisphere, associated with a higher surface

albedo and hence higher rates of photolysis, is another distinction between conditions determin-

ing concentrations of OH in the two hemispheres. The present distribution of OH results in

nearly equal hemispheric means for OH, mean tropical concentrations are slightly higher in the

south, whereas more OH is predicted for midlatitudes in the north (Table 4).

Lately two industrial compounds, C2Cl4 and CH2Cl2 , with relatively short lifetimes with

respect to reaction with OH (about 3 and 4.5 months in the tropics, respectively), have been

added to the list of species with documented releases and observational constraints. Observa-

tions of both gases [CMDL report #23, 1996] are compared with simulations in Figures 16 and

17; only industrial sources and the OH-sink are included in the simulations (unless noted). Esti-

mates of emissions are from McCulloch and Midgley [1996] for 1989-1992 and Midgley [1998,

pers. comm.] for 1993-1996. For C2Cl4 , the rate constant for reaction with OH is from DeMore

et al. [1997], whereas for CH2Cl2 we used 1.92.10−12e−897/T , which represents a fit to 4 recent

measurements, all within 15% of each other [DeMore, 1997, pers. comm.]. Simulations were

initialized in 1988 from the distribution observed by Koppmann et al. [1993]. Here we focus

solely on the overall levels of these compounds as a possible constraint for mean hemispheric

and semi-hemispheric concentrations of OH; we removed seasonal variations by considering

12-month running means.

For CH2Cl2 , model results for standard OH (bold solid lines) are ∼ 15-20% higher than

observations in the northern hemisphere and ∼ 10-15% lower than observations in the southern

hemisphere (Figure 16). For C2Cl4 , model results are higher than observations in the northern

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- 32 -

hemisphere (the discrepancy is larger than for CH2Cl2), and on average are close to observed

levels in the southern hemisphere (Figure 17). A comparison of the discrepancies for the two

tracers (or alternatively, a comparison with observations for the ratio of concentrations of the

two tracers, which is insensitive to the distribution of OH [Singh et al., 1996]) suggests that a

significant loss process for C2Cl4 , unrelated to OH, is missing in the simulation. Singh et al.

[1996] and Rudolph et al. [1996] used the high reactivity of C2Cl4 with Cl (300-400 times

higher than with OH) to define an upper limit for the abundance of Cl in the troposphere. A

simulation of C2Cl4 allowing for day-time-mean concentrations of Cl of 1.25×104mol cm−3 in

the lowest 500 m over the oceans (consistent with estimates of these authors), displays

discrepancies with observations similar to those for CH2Cl2 : concentrations are too high by ∼

20% in the northern hemisphere and too low in the southern hemisphere (see Figures 16 and 17).

Alternatively, a 30% increase in the rate constant for reaction of C2Cl4 with OH (at the upper

limit of the present uncertainty [DeMore et al., 1997]) leads to similar results (compare short-

and long-dashed lines in Figure 17).

For such tracers as CH2Cl2 and C2Cl4 , i.e., relatively short-lived and emitted mainly at

northern midlatitudes, a uniform scaling of the loss frequency globally (e.g., by scaling the rate

constant) affects disproportionally concentrations in the southern hemisphere. As expected for a

tracer with a nearly-steady-state behavior, the scaling of the loss frequency globally by -25% and

+50% leads to a change in its global abundance by about ±33% (the higher sensitivity towards

decrease follows from the inverse dependence on loss frequency). Concentrations at northern

midlatitudes display a lower than average sensitivity, with a response of less than +17 and -22%,

respectively, because they are determined to a large degree by a balance between the rates of

emissions and transport to the tropics. In contrast, at southern midlatitudes such modifications of

the loss frequency globally lead to a larger than average change in calculated concentrations of

about +59% and -56%, respectively, reflecting the cumulative effect of modified losses of tracer

en route from northern midlatitudes and thus in the flux of tracer available for the region to the

south. The current ±30% uncertainty in the rate constant for reaction of C2Cl4 with OH leads to

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- 33 -

the uncertainty in simulated annual mean concentrations of C2Cl4 at northern midlatitudes from

-13% to +22%; however, at southern midlatitudes, the same uncertainty is from -35% to +70%

as depicted by the distance between dotted and short-dashed lines in Figure 17. Of the two

tracers, CH2Cl2 and C2Cl4 , the former appears better suited at present for constraining OH,

given large uncertainties in the magnitude of its chlorine sink and in the rate constant for reac-

tion with OH.

Large modifications of the distribution of OH would be needed to eliminate the discrepan-

cies between model results for CH2Cl2 and observations, the excess of ∼ 15-20% in the northern

hemisphere and the deficit of 10-15% in the southern hemisphere (Figure 16), because: (1) the

sensitivity of concentrations of CH2Cl2 to an increase in OH in the northern hemisphere is

impeded by the compensating modification of the flux of CH2Cl2 out of the region, and (2) an

increase in OH in the north would further decrease concentrations of CH2Cl2 in the southern

hemisphere by decreasing the flux into the region. If we were to attribute the ∼ 20% excess of

CH2Cl2 in the northern hemisphere in the standard simulation solely to underestimated OH lev-

els, most of it would be eliminated by an increase in concentrations of OH in the northern hemi-

sphere by 35%; however, this modification would further diminish computed concentrations of

CH2Cl2 in the southern hemisphere. In order to compensate for this decrease in the southern

hemisphere as well as for the apparent deficit simulated in that region in the standard simulation,

concentrations of OH in the southern tropics would have to be decreased by about 60% (chain-

dashed lines) resulting in a decrease by ∼ 45% in the mean hemispheric concentration. Given

our current knowledge of distributions of precursors for OH and the limited sensitivity of con-

centrations of OH to their specification, it is hard to contemplate errors of that magnitude. As

will be clear from the summary given in Figure 13, independent observational constrains render

such large errors unlikely.

At the same time, other imperfections of the model, unrelated to OH, may be contributing

significantly to discrepancies between simulated and observed levels of CH2Cl2 . Concentrations

of such tracers as C2Cl4 and CH2Cl2 , particularly in the southern hemisphere, are more sensitive

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- 34 -

to the rate of interhemispheric mixing than those for the long-lived species, CFCs [Cunnold et

al., 1994] and 85Kr [Weiss et al., 1983] used to test the accuracy of this rate in the model

[Prather et al., 1987; Jacob et al., 1987]. The current CTM includes a parameterization of hor-

izontal mixing with the diffusion coefficient proportional to the intensity of local convection

and to the square of the length parameter D, introduced by Prather et al. [1987] to account for

processes unresolved in the parent GCM. In the standard CTM, D=180 km is used. Figure 18

illustrates small differences, as compared to observational constraints, in results for 85Kr with

mixing lengths of 180 km and 250 km. The simulation of CH2Cl2 with the standard OH, but

with D=250km, eliminates half of the discrepancy for CH2Cl2 in the northern hemisphere, and

produces an excess in the southern hemisphere similar to that remaining in the northern hemi-

sphere (thin solid lines in Figure 16). In reality the sensitivity of the distribution of CH2Cl2 to

representation of the north-south transport may even be greater than is shown in Figure 16, if

larger errors in the rate of mixing occur in the season when the interhemispheric gradient of

CH2Cl2 is at its peak (northern winter - spring) than in other seasons; seasonal variations in the

interhemispheric gradient of passive tracers are small and their observations do not constrain

sufficiently the seasonality of the interhemispheric mixing [Prather et al., 1987].

The simulation of CH2Cl2 with D=250 km results in concentrations of CH2Cl2 higher than

observed by 5-10% uniformly throughout the globe. Emissions of CH2Cl2 may be too high by

as much as 5% [Midgley and McCulloch 1996] and a 5% uncertainty is associated with the abso-

lute calibration of CH2Cl2 . Thus a globally-uniform discrepancy of 10% could be eliminated if

both calibration and emissions were assumed at the limit of their uncertainty.

Results for CH2Cl2 (expressed as annual means) suggest that the simulated mean concen-

tration of OH in the northern hemisphere is either accurate or too low by less than ∼ 26%) and in

the southern hemisphere is either accurate or too high by less than 82%. Crosses in Figure 13 for

CH2Cl2 correspond to the middle of these ranges. We emphasize however that the estimate of

the mean concentration of OH for the southern hemisphere consistent with observations of

CH2Cl2 depends strongly on that adopted for the northern hemisphere.

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- 35 -

Simulations discussed above suggest that industrial sources of CH2Cl2 are sufficient for

balancing the budget of CH2Cl2 for the present distribution of OH, using the newly measured

rate constant. It is apparent from Figure 16, that results for the standard OH and slightly

increased interhemispheric mixing are higher than observations throughout the globe (by ∼

10%). This increase in the interhemispheric mixing results in the decrease in the simulated glo-

bal abundance of CH2Cl2 of ∼ 1%. Non-industrial sources of CH2Cl2 , from the ocean and from

biomass burning, have been suggested in earlier studies [Singh et al., 1996, Rudolph et al., 1996]

using a higher rate constant. If however these sources are proved to represent a significant frac-

tion of the industrial source, it may be difficult to reconcile constraints for the global abundance

of OH imposed by long-term trends in CH3CCl3 and HCFC-22 with those given by the levels of

CH2Cl2 (see Figure 13 under "global mean"). An addition of 40 Gg year−1 of CH2Cl2 distri-

buted uniformly over the ocean surface (which amounts to only 13% of the industrial source

averaged over 1994-1996) results in concentrations of CH2Cl2 higher than observed in both

hemispheres, for standard OH and standard mixing (dotted lines in Figure 16).

6.2 Observations of CH3CCl3 as a constraint for the ratio of mean hemispheric concentrations

of OH. The present distribution of OH results in similar hemispheric means (integrated with

respect to mass of air), with the north to south ratio of 0.99 (Table 4). As follows from Section

6.1, annual means of CH2Cl2 constrain the ratio of its mean hemispheric loss frequencies to a

range from ∼ 1 to 2.4. Corrections of the computed distribution of OH resulting in the ratios

larger that 2 would be inconsistent with our present understanding of uncertainties in the calcu-

lated values (for the present chemical mechanism), and those in the range 1.5-2 would be

unlikely.

Observations of tracers with lifetimes significantly longer than the interhemispheric

exchange time, such as CH3CCl3 , may help constrain the ratio of mean hemispheric concentra-

tions of OH (integrated with respect to the mass of CH3CCl3 and weighted by the rate constant).

For example, in a steady state with a constant rate of emissions located in the northern hemi-

sphere, the net flux to the southern hemisphere could range from zero to the full emissions, if the

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- 36 -

loss were confined to one hemisphere, northern or southern, respectively; in the case of sym-

metric loss frequencies, the flux would be equal to half of emissions. The interhemispheric gra-

dient would adjust accordingly to ensure the transmission of a proper flux. The interhemispheric

gradient of CH3CCl3 is however sensitive to the rate of interhemispheric mixing in the model,

e.g., if interhemispheric transport is too vigorous, an erroneous interhemispheric bias, with

higher loss frequencies in the southern hemisphere, would be deduced from a comparison of the

simulated and observed gradients of CH3CCl3 . For CH3CCl3 , observations of 85Kr [Weiss et

al., 1983] and CFCs [Cunnold et al., 1994] should provide a sufficiently accurate test of

interhemispheric mixing (see Figures 18 and 19) because the lifetime of CH3CCl3 exceeds the

interhemispheric exchange time of ∼ 1 year [Prather et al., 1987; Jacob et al., 1987] by about a

factor of 5 (in contrast to CH2CCl2, with lifetime of several months, as discussed in Section

6.1). In Figure 20, the interhemispheric gradient of CH3CCl3 is compared with observations

[Prinn et al., 1995] for simulations using the standard OH as well as for distributions of OH

obtained by scaling concentrations in each hemisphere to give the same global mean lifetime of

CH3CCl3 , but with the north-south ratios of hemispheric means for OH of 2 and 0.5. We chose

to average the latitudinal distribution over 1980-1983 because during this time, emissions were

relatively stable, within 3% of the mean. (Modifications of the interhemispheric gradient associ-

ated with the El Nino from mid-1982 to mid-1983 [Prinn et al., 1992], are less than 10%, and

may affect the mean gradient over 1980-1983 by a few percent at most.)

The location of the ITCZ in the model, as the actual boundary between the hemispheres,

affects the interhemispheric gradient of CH3CCl3 in two ways: (1) by influencing the effective

mean concentration of OH in each hemisphere (since the near-equatorial region in question is

characterized by particularly high concentrations of OH), and (2) by defining the effective dilu-

tion volume in the northern hemisphere. The simulation of 85Kr reproduces the observed latitu-

dinal distribution in boreal winter, spring and autumn (as shown in Figure 18 for March) giving

no indication of a significant misplacement of the ITCZ over the Atlantic. However, a com-

parison of observed and simulated peak intensities of the Hadley circulation [Rind and Lerner,

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- 37 -

1996] suggests that on average the ITCZ in the model may be shifted north by one grid box in

latitude, i.e., by 4°. This would result in an exaggeration of the interhemispheric gradient of

CH3CCl3 , leading erroneously to an underestimation of the north-to-south ratio for OH. (The

shift in the location of the ITCZ to the north may be partly responsible for the overestimation of

the interhemispheric gradient of CH2Cl2 discussed in Section 6.1.) For the present distribution

of OH, a 4° latitude shift of the ITCZ (for the full range in longitude) results in less than 10%

difference in the effective ratio of interhemispheric means for the loss frequency of CH3CCl3 ,

which would lead to a less than 5% error in the interhemispheric gradient of CH3CCl3 , as

estimated using a 2-box model. The effect of the decrease in the dilution volume associated with

the possible shift of the ITCZ is not evident in results for 85Kr (Figure 18).

The present distribution of OH, with nearly equal hemispheric means, provides the best

agreement with observations, whereas depletion of OH by a factor of 2 or more in one hemi-

sphere as compared to the other is inconsistent with observations. For a given global mean OH,

constraining the ratio of hemispheric means to a factor of 2 is equivalent to constraining hem-

ispheric mean concentrations of OH to better than ±33%. No ocean sink was included in these

simulations. Assuming that the ocean sink amounts to 4-8% of the total sink of CH3CCl3 , and

that it is proportional to the area of the ocean, we conclude that the omission of the ocean sink in

the model may introduce a bias of a few percent in estimates of hemispheric means for OH

(underestimation of OH in the north and overestimation in the south).

It is difficult to tighten significantly the range of this estimate because of uncertainties asso-

ciated with the definition of simulated "background levels" of CH3CCl3 as well as CFCs at the

northern hemisphere sites, in the proximity of the source regions. For example, the simulation

with standard OH underestimates concentrations at Barbados (13°N), and overestimates them at

Cape Meares (45°N) (Figure 20). Results for CFC-12 (not shown) display similar errors at Bar-

bados and Cape Meares, but not for CFC-11 (Figure 19). We attribute these discrepancies to

errors in representation of short-range transport of tracers from source regions, to uncertainties in

the distribution of emissions within large countries and to difficulties of relating grid-box results

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- 38 -

to observations at a site in the region of steep gradients.

In the next several years, when emissions of CH3CCl3 will nearly cease as envisaged in the

Montreal Protocol, observations of CH3CCl3 may present a unique opportunity to improve the

constraint on the ratio of hemispheric means for OH as discussed in Section 8. We emphasize

that the interhemispheric gradient of CH3CCl3 is insensitive to decreases (as well as moderate

increases, within 25%) in its global mean loss frequency if the latitudinal distribution is kept the

same [S90; Spivakovsky, 1991].

6.3 Seasonal variations as a constraint for OH at temperate latitudes. The relative ampli-

tude of seasonal variations of CH3CCl3 observed at Tasmania (41°S) contains a measure of the

tropospheric concentration of OH integrated over southern extratropics (weighted by frequency

of reaction with CH3CCl3) [S90]. Figure 21 (solid line) compares relative amplitudes of simu-

lated and observed seasonal variations of CH3CCl3 averaged over 1980-1991 at Tasmania

(41°S). Dashed lines represent results for the standard OH reduced and increased by 50% south

of 28°S. This comparison suggest that computed concentrations of OH in that region may be too

high by ∼ 15-25%. Results for OH increased by 50% are outside of one σ of observations.

An uncertainty of about 25% is associated with the magnitude of the dynamical component

of seasonal variations at Tasmania for long-lived tracers emitted mainly at northern midlatitudes

(thin solid line in Figure 21). The present record for CFCs at Tasmania spans over 17 years and

includes 3 sets of observations [Cunnold at al., 1994], none of which exhibits a statistically

significant annual cycle (at the confidence level of 95%). The CTM simulation of CFCs predicts

small seasonal variations at Tasmania, with amplitudes less than 0.3%, i.e., comparable to the

instrumental noise in observations [Cunnold et al., 1994]. The present model uses meteorologi-

cal fields from one year of the GCM simulation therefore model results characterize seasonal

variations in transport over the course of one year rather than a recurring annual cycle.

Similarly to results for CH3CCl3 , a comparison of simulated and observed seasonal varia-

tions for CH2Cl2 at Tasmania suggests an overestimation of OH in that region, but of somewhat

larger magnitude, by 25-50% (Figure 22). Opposite to conclusions for CH3CCl3 and CH2Cl2 ,

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- 39 -

results for C2Cl4 (Figure 23) suggest that its loss frequency in summer should be increased by

40-50%, consistent with an increase of the rate constant and/or the presence of the chlorine sink

discussed in Section 6.1. If the chlorine sink determines a significant part of the disparity of the

results for CH2Cl2 and C2Cl4 evident in Figures 22 1nd 23, its seasonality must be similar to

that of the OH sink (Figure 23).

In northern midlatitudes, observations of a variety of industrial pollutants with relatively

short lifetimes can potentially be used to evaluate computed concentrations of OH in that region.

However, uncertainties in the magnitude and distribution of sources present a major difficulty.

In addition, the proximity of source regions to sampling locations contributes greatly to the vari-

ability in observations and complicates interpretation of measurements using models with a hor-

izontal resolution typically of hundreds of kilometers across a grid box. Goldstein et al. [1995a]

circumvented some of these difficulties by defining "background concentrations" at Harvard

Forest (42.5°N, 72.2°W) possible because of continuous, high-frequency observations. Using

the relative amplitude of seasonal variations, they demonstrated that the seasonal behavior of

ethane, acetylene, propane, butane, pentane and hexane, with lifetimes in summer ranging from

40 to 1.5 days, can be explained in the context of a simple 1-box model, with constant rate of

emissions. Their work suggests that seasonal variations in OH in fact determine the seasonality

of background levels of these compounds at the site, as is predicted by simulations using the

CTM. Mean tropospheric concentrations of OH presented here are higher by 20% north of 32°N

than those used by Goldstein et al. [1995a] (see Table 4), however, the currently recommended

rate constant for reaction of C2H6 with OH is lower by about 15% [DeMore et al., 1997]; in

addition, less OH resides in the lower troposphere in the present distribution as discussed in Sec-

tion 2. Consequently, our simulation of seasonal variations of C2H6 at Harvard Forrest (Figure

24) leads to results similar to those of Goldstein et al. [1995a]: average concentrations of OH at

northern midlatitudes are constrained to ±50%, with the simulation for standard OH falling near

the middle of the range consistent with observations.

For CH2Cl2 , the magnitude of seasonal variations in northern extratropics is consistent with

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- 40 -

observations for the simulation with standard OH, and is outside the range of observations for

the simulation with OH increased by 50% (Figure 25). For C2Cl4 (Figure 26), results suggest

that the loss frequency using standard OH is too low. A comparison of the results for C2Cl4

with those for CH2Cl2 (Figures 25 and 26) is consistent with results for annual means

throughout the globe (Figures 16 and 17) and for seasonal variations at southern midlatitudes

(Figures 22 and 23), suggesting that a Cl sink of moderate strength (within the range estimated

by Singh et al. [1996] and Rudolph et al. 1996]) and with the seasonality similar to that for OH

sink is missing in the simulation of C2Cl4 or, alternatively, that the rate constant for reaction

with OH is too low.

Unlike other constraints, those afforded by the relative amplitude of seasonal variations are

insensitive to uncertainties in the absolute calibration and magnitude of the industrial emissions;

they are sensitive however to errors in representing the seasonality of emissions and transport.

7. The utility and limitations of observations of 14CO as

a test for OH.

Since the pioneering work of Weinstock and Niki [1972], 14CO has been considered as one

of the few gases that can provide a measure of tropospheric abundance of OH, because the mag-

nitude and distribution of its source, mainly of cosmic origin, were believed to be known, and

reaction with OH constituted the major sink. By virtue of its relatively short lifetime (1-1.5

months in the tropics and midlatitude summer), observations of 14CO were expected to provide

estimates of OH on the hemispheric and semi-hemispheric scale [e.g., Volz et al., 1981; Bren-

ninkmeijer et al., 1992].

However, interpretation of observations of 14CO proved to be difficult because it involves

simulation of complicated dynamical processes which cannot be readily tested. About 50% of

the cosmic production of 14CO outside the tropics occurs above 150 mb. Therefore the effective

tropospheric source of 14CO (of cosmic origin) in in models may differ by as much as a factor of

2 in the extratropics, depending on the flux from the stratosphere which in turn is determined by

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- 41 -

the relative magnitudes of the stratospheric loss frequency and the rate of transport to the tropo-

sphere. Results may appear insensitive to the rate of troposphere-stratosphere exchange [Mak et

al., 1992] if loss frequencies in the stratosphere are too low. Recent successes in measuring and

modeling concentrations of OH in the lower stratosphere [Wennberg et al., 1994] may reduce

significantly uncertainties in the rate of chemical destruction of 14CO in the stratosphere. How-

ever, adequate testing of the rate of transport from the stratosphere remains problematic; obser-

vations of excess 14CO2 after nuclear bomb tests [Johnston, 1989] customarily used for that pur-

pose, although helpful for identifying gross errors, do not allow discrimination between (1)

erroneous rates for the strat-trop exchange and (2) inaccurate representation of transport within

the stratosphere, north to south and to higher altitudes [Shia et al., 1989; Jackman et al., 1991;

Prather and Remsberg, 1993].

Another factor affecting simulated tropospheric concentrations of 14CO, particularly out-

side the tropics, is the height of the tropopause in the model, especially at high latitudes. Mean

production rates of cosmic 14CO outside the tropics between 250 and 200 mb and between 200

and 150 mb amount to about 35% and 45%, respectively, of the production rate below 250 mb.

If the model tropopause is too high, a region of the stratosphere with fast production is errone-

ously attributed to the troposphere, and 14CO produced in this region is allowed to escape des-

truction by stratospheric OH. This leads to an overestimation of the tropospheric source of

14CO. (Conversely, if the tropopause in the model is too low, tropospheric concentrations may

be significantly underestimated.) Most of the models used thus far to interpret observations of

14CO, including the present model, were developed with an emphasis on tropospheric processes

and lacked both adequate resolution in the tropopause region, and the appropriate physics [cf

Spivakovsky and Balkanski, 1994, Mak et al., 1994]

Equally important is representation of intrahemispheric transport in the troposphere (on a

seasonal time-scale) because most of 14CO originates at high and middle latitudes, and the

highest loss frequencies occur in the tropics. More vigorous mixing between midlatitudes and

the tropics would decrease the intrahemispheric gradient and expose higher levels of 14CO to

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- 42 -

tropical loss frequencies, resulting in a lower tropospheric abundance of 14CO. In the northern

hemisphere, observations of CFCs and 85Kr provide some constraint for the rate of intrahem-

ispheric mixing, but in the southern hemisphere they display little latitudinal gradient and thus

provide only a lower limit for this rate (Figure 18).

Because of the difficulties discussed above, observations of 14CO at midlatitudes cannot be

regarded as a definitive measure of OH. Nevertheless, they may help identify model defects, and

in conjunction with other tracers provide useful constraints as illustrated below.

The distribution of cosmic emissions and their dependence on the sunspot number in our

simulations of 14CO were taken from Lingenfelter [1963] and O’Brien [1979], respectively,

assuming that 95% of 14C is emitted as 14CO [Volz et al., 1981]. Non-cosmic emissions of

14CO (biomass burning and oxidation of CH4 , isoprene and other NMHC) were computed using

emissions of CO from Wang et al. [1998] and 14CO/12CO ratios in individual sources from

Volz et al. [1981]. As shown in Figure 27a, the standard model (solid line) significantly overesti-

mates the relative amplitude of seasonal variations of 14CO at southern midlatitudes which in the

model is determined by seasonality in OH. This may suggest [S90, Goldstein et al., 1995a], that

computed concentrations of OH are too high in that region. However, the model also overesti-

mates the annually averaged concentration (as can be seen in the lower panel of Figure 27b

showing the same data and simulations as in Figure 27a, but on the absolute scale, with the

implication that OH levels may be too low. The apparent contradiction indicates an inadequate

representation of the effective net flux of 14CO reaching the lower troposphere at southern

midlatitudes, either in its magnitude, or seasonality, or both.

In particular, a poor representation of the stratosphere may explain a large part of the

discrepancy. The stratosphere in this model is resolved in two layers in the extratropics, from

150 to 70 mb, and from 70 to 10 mb. The tropopause in the model is located at about 150 mb

outside the tropics all year around, which is too high except in summer [e.g., Holton et al.,

1995]. A simulation of the evolution of the global distribution of 14CO2 after the nuclear bomb

tests [Johnston, 1989] suggests that the rate of transport from the stratosphere in the model may

-- --

- 43 -

be too high by as much as a factor of 2 (see Figure 28). This rate may be particularly excessive

in the southern hemisphere, where it is slightly higher than in the north, in contradiction to the

analysis of Rosenlof and Holton [1993], Yang and Tang [1996] and Eluszkiewicz et al. [1996]

indicating more vigorous strat-strop exchange rates in the north. (Observations of bomb 14CO2

are available only at 42°S in the southern hemisphere, and do not extend above 22 km.)

To estimate the magnitude of the error in our simulation of 14CO due to inadequate

representation of the stratosphere, we (1) reduced the air-flux across the 150 mb surface by a fac-

tor of 2 based on the results for bomb 14CO2 , and (2) moved the portion of stratospheric emis-

sions that the model erroneously attributes to the troposphere (because of the fixed position of

the model tropopause) to the model layer above 150 mb; the seasonal tropopause heights used to

estimate the magnitude of the misplaced emissions were taken from McCormick et al. [1993].

As can be seen in Figures 27, 29 and 30, in austral winter in particular, simulations with the

reduced flux from the stratosphere still exaggerate not only the absolute concentrations of 14CO

but also its vertical and horizontal gradients in the southern hemisphere. Our analysis of GCM II

temperature fields revealed highly exaggerated horizontal gradients and uncharacteristically low

temperatures poleward of 60°S below 400 mb in winter; e.g., at 600 mb, the temperature gra-

dient between 60°S and 80°S is about 20°C higher than observed: 28° in the GCM vs. 8° in

Peixoto and Oort [1992]. In addition, the jet stream in the southern hemisphere is about twice as

strong as observed, whereas the transient eddy energy is lower by more than a factor of two than

observed, with a maximum displaced towards the equator by more than 20°. These discrepan-

cies in the GISS GCM II dynamics may be an indication of insufficient mixing between southern

high latitudes and the tropics [cf Rind and Lerner, 1996]. Unfortunately, this is one of the

aspects of transport that is not readily testable. The latitudinal distribution of 85Kr changes little

if an additional arbitrary horizontal diffusion of 6.1010cm2sec−1 south of 28°S is introduced

below 400 mb in austral winter, combined with vertical diffusion of 8.105cm2sec−1 south of

60°S (Figure 18, the curve is indistinguishable from that for standard interhemispheric mixing,

i.e., for D=180 km). This modification of model dynamics sufficiently decreases the absolute

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- 44 -

level of 14CO in the southern hemisphere, as well as its gradient and the relative amplitude of its

seasonal cycle (Figures 27, 29, and 30). Thus, the sensitivity of extratropical levels of 14CO to

the largely untestable aspects of model dynamics limits the utility of observed absolute concen-

trations of 14CO as a measure of OH in that region, although they provide a stringent test of an

overall performance of the model in simulating combined effects of chemistry and transport.

The simulated annual cycle of 14CO at northern midlatitudes is compared with observa-

tions in Figure 31, relatively to the annual mean (a) and on the absolute scale (b). The relative

amplitude for the simulation with standard OH is consistent with observations, while the abso-

lute concentrations are low even for the standard simulation (solid lines), i.e., for the uncorrected

flux from the stratosphere. Therefore, results for 14CO at northern midlatitudes suggest that a

significant underestimation of concentrations of OH in the northern extratropics is unlikely.

As can be seen in Figure 29, in the tropics, concentrations of 14CO are less sensitive to

modifications of the flux from the stratosphere, or to the rate of intrahemispheric mixing in the

troposphere because about two-thirds of 14CO in the tropical troposphere is produced in the tro-

pics, with about a half of that amount coming from biomass burning and oxidation of hydrocar-

bons. Using observations of 14CO in the tropics, Mak et al. [1992] first suggested that concen-

trations of OH from S90 were too low. Results for the present distribution are consistent with

observations of 14CO [Mak et al. 1992; 1994] at 10°S- 30°S in January-February and at 0°-30°N

in August; however, in winter in the tropics of both hemispheres, simulated concentrations of

14CO are too high by 10-15% with an implication that concentrations of OH are too low in that

season by 10-15%. While in the northern tropics, results for annual means of CH2Cl2 also sug-

gested that concentrations of OH may be too low, in the southern tropics the opposite conclusion

was implied by observations of CH2Cl2 (see Figure 13). Because of the significance of the

non-cosmic sources of 14CO in the tropics, a simulation including all CO isotopes in the OH-CO

coupled model would be particularly beneficial for testing both sources of CO and the computa-

tion of OH.

8. Last offerings of a departing friend.

-- --

- 45 -

Consumption of CH3CCl3 , defined as "production + imports - exports", was to cease by January

1996 in all but developing countries (Article 5 of the Montreal Protocol). Emissions are

expected to diminish sharply as reserves in the developed world are exhausted [P. Midgley, pers.

comm.]. The developing countries are allowed to consume up to 100 Gg y−1 until 2010. How-

ever, in 1993 and 1994-1996, total emissions in developing countries were only about 51 Gg y−1

and 30 Gg y−1 , respectively, and total emissions in the following years are not expected to

exceed significantly 30 Gg y−1 [Midgley and McCulloch, 1995; P. Midgley, pers. comm.], i.e.,

less than 6% of the mean rate of emissions in 1980-1983. Were the emissions of CH3CCl3 to

cease entirely, the rate of relative decrease of the global abundance of CH3CCl3 would unambi-

guously define the lifetime of CH3CCl3 , free of the uncertainties in the absolute calibration of

measurements and the magnitude of emissions, associated with the current estimate [Ravishan-

kara and Albritton, 1995]. Even if emissions were to linger near ∼ 30 Gg y−1 , the impact of

these uncertainties will greatly diminish during the next 3 years because atmospheric levels of

CH3CCl3 will be at least factor of 4 higher than the steady state at 30 Gg y−1 (about 6 ppt).

As another benefit of diminished emissions, we expect that observations of CH3CCl3 over

the next several years will allow us to tighten constraints for the loss frequency of CH3CCl3 on

scales smaller than global. In previous decades, seasonal variations of CH3CCl3 at the surface at

northern midlatitudes were determined largely by the seasonality of convection and other

processes dispersing CH3CCl3 away from the source regions. Under the new conditions, an

annual cycle of CH3CCl3 at northern midlatitudes is expected to bear mainly the signature of

seasonal variations in the loss frequency, without significant interference from short-term varia-

bility ("pollution events"). As can be seen in Figures 32 and 33, for simulations discussed

below, with no emissions or with small lingering emissions, zonally averaged results (lines)

differ little from results at ALE/GAGE and NOAA sites (circles), in sharp contrast to the mani-

fest zonal asymmetry simulated for the distribution of CH3CCl3 in the 1980s [S90]. Improve-

ment of accuracy is expected for southern midlatitudes as well. In previous decades, the distri-

bution of CH3CCl3 was characterized by a positive north-south gradient, including that between

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- 46 -

the southern tropics and extratropics, maintained by the northern hemisphere emissions. With

emissions at present levels and lower, it is expected that a negative gradient between the tropics

and southern midlatitudes will become a permanent feature of the distribution of CH3CCl3 from

now on as the global distribution of CH3CCl3 becomes determined more by the latitudinal distri-

bution of its loss frequency rather than that of the emissions. A comparison of observed seasonal

variations of CH3CCl3 at southern midlatitudes for the two periods, characterized by the oppo-

site signs of the gradient, may help to determine the role of the dynamical component and thus

tighten constraints for OH levels in that region.

The interhemispheric gradient of CH3CCl3 has decreased dramatically in the last few years

[e.g., Prinn et al., 1995; Montzka et al., 1996; Th. Thompson, pers. comm, 1998] as a result of

diminishing releases [P. Midgley, pers. comm., 1997]. In the case of zero emissions, a

significant interhemispheric asymmetry in the loss frequency of CH3CCl3 would manifest itself

in an interhemispheric gradient that, when averaged over the year, will differ significantly from

zero, as shown in Figure 32 (left panels). Here a simulation with the standard OH (solid line) is

compared to those for the north/south ratios of hemispheric means of 1.5 and 1/1.5, but with the

same global mean concentration of OH. Results at the surface are shown as a difference with

those at 42°S, relative to the global average of surface concentrations. Simulations were initial-

ized from observations at the end of 1996 [Th. Thompson, pers. comm, 1997]. At that time, the

global mean concentration of CH3CCl3 was about 90 ppt, with a gradient between northern and

southern high latitudes of about 9 ppt. Results are shown for 1999 and 2002; as expected, on the

relative scale, they are virtually the same for different years. A factor of 1.5 bias in the mean

hemispheric loss frequencies of CH3CCl3 leads to a gradient between midlatitudes of the two

hemispheres of 4% (averaged annually). For comparison, the current precision of ALE/GAGE

observations is better than 0.5% [Prinn et al., 1995]. For standard OH, the north-south gradient

fluctuates from about +3.5% in February to -3.5% in August (Figure 33, left panels). In contrast,

it remains positive throughout the year for the simulation with OH depleted in the southern hem-

isphere (ranging from about 7% in northern winter to nearly zero in summer), and negative

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- 47 -

throughout the year for the opposite case. The absolute amplitude of seasonal variations of the

gradient gradually subsides in proportion to the atmospheric level of CH3CCl3 , and hence,

results on the relative scale differ little from year to year.

Right panels in Figures 32 and 33 show results for a constant rate of emissions of 30 Gg y−1

(chosen arbitrarily) from 1998 on, and distributed over developing countries only, in accordance

with Midgley and McCulloch [1995]. This distribution results in about 10% of world emissions

in the southern hemisphere (as compared to only 3% in 1992 and 6% in 1993). In contrast to

simulations with zero emissions (in which the interhemispheric gradient was maintained solely

by the difference in loss frequencies in the two hemispheres), results for different years vary on

the relative scale because the part of the gradient determined by emissions remains constant

while atmospheric levels of CH3CCl3 decline. If emissions are monitored, the relative observa-

tions of the latitudinal distribution of CH3CCl3 , conducted with carefully maintained precision,

could allow one to constrain the interhemispheric ratio of its loss frequency to a factor of 1.5 or

better, while the global burden remains high (compared to the steady state corresponding to the

level of lingering emissions). This would be equivalent to constraining hemispheric mean loss

frequencies of CH3CCl3 to ±20%, assuming that the global mean is known. If it will not be pos-

sible to monitor emissions, then observations of the latitudinal distribution of CH3CCl3 over the

next several years may prove useful if concentrations of CH3CCl3 turn out to be significantly

lower on average in the northern than in the southern hemisphere. This would signify

significantly higher levels of OH in the northern hemisphere than in the south, since the major

part of the emissions is expected to occur in the northern hemisphere.

The lifetime of CH3CCl3 in the troposphere (computed using the present distribution of

OH) varies in the range 3-3.5 years between 20°S and 20°N, and then increases sharply with lati-

tude to values greater than 10 years at 60° (Figure 34). In the absence of latitudinal mixing, con-

centrations would decrease by 25-30% per year at low latitudes and by less than 10% at high

latitudes, resulting in the depth of the tropical dip of 15-20%. The predicted depth however is

about 4%, i.e., about a factor of 4-5 less, suggesting an exponential decay with the rate of about

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- 48 -

7 months for the gradient between the zone of fastest loss and high latitudes. Thus, the latitudi-

nal distribution of CH3CCl3 , in the absence of a significant interference from pollution events

will also present a test of the combined effects of (1) disparity between loss frequencies in the

tropics and extratropics and (2) intrahemispheric mixing.

9. Conclusions.

The climatological tropospheric distribution of OH presented here results in a global annual

average of 1.16.106 mol cm−3 (integrated with respect to mass of air up to 100 mb within 32°

latitude and up to 200 mb outside of that region). The predicted global average concentration of

OH is unlikely to change by more than 10% due to improvement in the specification of any

given precursor. A larger change may occur if significant modifications for several precursors

were to affect concentrations of OH in the same direction. There is no reason however to expect

that errors in distributions of various precursors should be correlated.

Measurements affording better definition of concentrations of O3 over the less polluted

regions, particularly in the tropics, and those for NOx in the region from 800 to 500 mb, would

be most effective in improving the accuracy of the computed OH. Uncertainties in the vertical

extent of isoprene levels in excess of 75-100 ppt may lead to errors in the global mean concen-

tration of OH on the order of 5%. For the present distribution, if isoprene chemistry were not

included, the global mean concentration of OH would increase by 4%. Inclusion of NMHC,

apart from isoprene, decreased the global mean OH by 7%.

While the global mean OH increased by 33% compared to that from S90, mean loss fre-

quencies of such tracers as CH3CCl3 and CH4 increased by only 23% because of redistribution

of OH within the tropospheric column: the fraction of OH residing below 700 mb in the present

distribution is lower than in S90 by 21%. This upward shift in OH occurred in part because of

the new distribution of clouds, and in part because of inclusion of NMHC combined with a

decrease in CO over most of the globe. As a result of the decrease in the fraction of the abun-

dance of OH residing in the lower troposphere, the value 277K used for scaling lifetimes of

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- 49 -

HCFCs and other long-lived gases (with a positive exponential dependence of rate constants on

temperature) to the known lifetime [Prather and Spivakovsky, 1990] has to be revised to 272K.

On average, there is little interhemispheric bias in the computed distribution of OH: the

ratio of mean hemispheric concentrations of OH, north to south, is 0.99. Mean tropical concen-

trations (within 32°) are 6% lower in the north, whereas 11% more OH is predicted for the

region poleward of 32° in the north than in the south.

The present distribution of OH results in a global annual mean lifetime of CH3CCl3 of 4.7

years [cf Prinn et al., 1995], including the stratospheric sink with the atmospheric lifetime of 43

years and allowing for the ocean sink of 78 years. As shown in Figure 13, the global mean con-

centration of OH (weighted by the loss frequency in reaction with CH3CCl3) is 3±3% higher

than is implied by the observed long-term trend in CH3CCl3 , using the most likely values for

non-OH sinks, absolute calibration, emissions and rate constant (the ±3% uncertainty is associ-

ated with the comparison of model results for CH3CCl3 with observations). An uncertainty in

determining mean OH implied by observations is estimated to be in the range from -24% to

+28% assuming that errors in all relevant values may affect the estimate in the same direction.

For HCFC-22, the present distribution of OH results in the global annual mean of 11.4

years, including the stratospheric sink of 229 years. The evolution of the global abundance of

HCFC-22 implies that the computed annual global mean concentration of OH (weighted by the

loss frequency in reaction with HCFC-22) is too high by 2±5%, with an uncertainty in the value

implied by observations estimated between -20% and +35%. Thus, results using long-term

trends of CH3CCl3 and HCFC-22 are consistent within a few percent (see Figure 13). About

80% of the loss of these compounds occurs in the tropics, with the similar dependence of loss

frequency on temperature. Only reported industrial emissions were used for HCFC-22.

Although significant uncertainties are associated with determining "true" global mean OH using

observed long-term trends of CH3CCl3 and HCFC-22 (dashed lines in Figure 13), consistent

results for two independent constraints afford a degree of confidence in these estimates.

Observed annually averaged levels of CH2Cl2 imply that the global annual mean for the

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- 50 -

present distribution of OH (weighted by the loss frequency in reaction with CH2Cl2) is too low

by 10±5%, whereas an uncertainty associated with defining the value for "true OH" is estimated

to be in the range from -20% to +30%. Part of the difference between the conclusion based on

annual means of CH2Cl2 and that based on long-term trends of CH3CCl3 or HCFC-22 may be

attributed to different weighting, because of a shorter lifetime (therefore steeper gradients) and

weaker dependence of loss frequency on temperature for CH2Cl2 than for CH3CCl3 or HCFC-

22. Industrial sources of CH2Cl2 are sufficient for balancing the budget of CH2Cl2 using the

recently measured rate constant and the present distribution of OH. Uncertainties in the strength

of the Cl-sink for C2Cl4 and in the rate constant for reaction with OH restrict at present the util-

ity of observations of C2Cl4 as a test of OH.

The sensitivity of tracer concentrations to rates of transport is an intrinsic difficulty in con-

straining the regional levels of OH, as reflected in large uncertainties associated with determin-

ing the relation between the "true OH" and that used in simulations (solid lines in Figure 13)

Using independent constraints based on observations of tracers with different lifetimes and dis-

tinct distributions of sources is essential for obtaining meaningful estimates of the accuracy of

computed OH on scales smaller than global. On the hemispheric scale, the rate of interhem-

ispheric mixing in the CTM affects greatly the conclusions drawn from observations of CH2Cl2 .

If this rate is taken at the upper limit consistent with observations of CFCs and 85Kr, observed

annual means of CH2Cl2 suggest that mean concentrations of OH may be accurate in both hemi-

spheres. If however the rate of mixing is taken at its lower limit, these observations indicate that

concentrations of OH in the north are too low by 26% and too high in the south by 82% and that

the ratio of hemispheric means for OH, north to south, is ∼ 2.5. The latter scenario contradicts

the evidence contained in the observed interhemispheric gradient of CH3CCl3 in the 1980s,

which rules out ratios greater than 2. Furthermore, other constraints do not support such large

corrections for concentrations of OH in either hemisphere.

In northern extratropics, tests using the relative amplitude of seasonal variations of

CH2Cl2 , C2H6 and 14CO do not suggest significant errors in the average levels of OH in that

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- 51 -

region (see Figure 13). Absolute levels of 14CO predicted for that region are too low even

though the flux of 14CO from the stratosphere in the model is too high, with the implication that

concentrations of OH at northern midlatitudes are not underestimated. Observations of 14CO in

the northern tropics in summer indicate accurate concentrations of OH computed for that region;

in winter however, concentrations of OH between the equator and 15°N appear too low by 20-

30%. Given all available evidence, some increase may be in order for computed concentrations

of OH in winter at low northern latitudes which may result in a small increase in the annual

mean concentration of OH for the whole northern hemisphere (5-10%). A greater underestima-

tion of OH in the northern hemisphere is unlikely because, in order to satisfy global constraints,

such as lifetimes of CH3CCl3 and HCFC-22, it would have to be compensated by a comparable

overestimation in the southern hemisphere. However, observations of 14CO in the southern tro-

pics call for an increase of concentrations of OH in that region in winter by 15-20%, and indi-

cate no deviation in summer. A decrease in OH by 10-15% in the southern hemisphere may be

indicated in the extratropics by the tests using the relative amplitude of seasonal variations of

CH2Cl2 , CH3CCl3 and 14CO (see Figure 13), but such a decrease would affect little the mean

concentration of OH for the southern hemisphere. One reason for underpredicting concentra-

tions of OH at low latitudes in winter of both hemispheres may be associated with uncertainties

in concentrations of O3 and NOx over the ocean during the biomass burning season. Observa-

tions of seasonal variations of CH3CCl3 , CH2Cl2 , 14CO and C2H6 offer no evidence for higher

levels of OH in the southern than in northern extratropics.

We used the interhemispheric gradient of CH3CCl3 as a constraint for the ratio of mean

hemispheric concentrations of OH. We emphasize however that the interhemispheric gradient of

CH3CCl3 is insensitive to any increases as well as to moderate decreases in the lifetime of

CH3CCl3 if the latitudinal distribution of its loss frequency remains unchanged.

Interpretation of observations of 14CO presents an additional level of difficulty as com-

pared to such tracers as CH2Cl2 because it requires an accurate simulation of the rate of tran-

sport from the stratosphere, of the stratospheric loss, and of the position of the tropopause; in

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- 52 -

addition, steep gradients between tropics and midlatitudes in both hemispheres, enhanced by a

shorter lifetime (about 1 month in the tropics) and high-latitude sources, make results at midlati-

tudes particularly sensitive to the rate of intrahemispheric mixing, which in the southern hemi-

sphere is not readily testable. Therefore, unless these other aspects of the model, unrelated to

OH, are proved accurate, absolute levels of 14CO outside the tropics cannot be regarded as an

unambiguous measure of the abundance of OH but rather as an important test of an overall per-

formance of CTMs.

Unless a significant change in the mechanism is recommended, it is difficult to contemplate

errors in excess of ±15% in global, hemispheric, and semihemispheric climatological averages

for the abundance of OH, given our present knowledge of the distributions of precursors and the

limited sensitivity of computed concentrations of OH to changes in their specification. It is

difficult however to test the computed values to that level of accuracy, especially on scales

smaller than the global.

In the next few years, while the global burden of CH3CCl3 remains high but emissions are

reduced, observations of CH3CCl3 will present a unique opportunity to constrain better global

and regional abundances of OH. The rate of change in the global burden of CH3CCl3 will be

less sensitive to errors in the absolute calibration or magnitude of emissions [Ravishankara and

Albritton, 1995]. A factor of 1.5 excess in the loss frequency in one hemisphere (as compared to

the other) would result in about a 4% gradient in the annual mean concentration between midlati-

tudes of the two hemispheres (with respect to the global mean at the surface). Interference from

"pollution events" in the next 3-4 years is not expected to confound the analysis of the latitudinal

gradient of CH3CCl3 in the next 3-4 years in a manner described above for the 1980s because

the global burden of CH3CCl3 in 1999-2002 will be much higher than the one in balance with

the rate of emissions, whereas in the 1980s, the imbalance was of the opposite sign. Our analysis

suggests that if emissions are monitored, observations of the latitudinal distribution of CH3CCl3

could allow us to constrain the interhemispheric ratio of its loss frequency to better than a factor

of 1.5 (equivalent to estimating mean hemispheric loss frequencies to ±20% assuming that the

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- 53 -

global mean is known). As in the past, most of the remaining emissions are expected to occur in

the northern hemisphere. If significant uncertainties in the magnitude of emissions cannot be

eliminated, the interhemispheric gradient of CH3CCl3 would provide an upper limit for the ratio,

north to south, of its mean hemispheric loss frequencies.

The estimates of the abundance of OH in the extratropics will be also improved under the

new conditions. In the north, both the interference of short-term variability and the seasonality

of transport rates are expected to have little impact on the annual cycle of CH3CCl3 as long as

the global burden remains high as compared to the steady-state level associated with the rate of

emissions. In the south, the negative north-south gradient between the tropics and midlatitudes

is expected to become a steady feature of the distribution of CH3CCl3 beginning in the summer

of 1997, as opposed to the positive gradient characteristic of previous decades. A comparison

between annual cycles of CH3CCl3 for the two periods, with opposite signs of the gradient, will

help constrain the magnitude of the dynamical component of seasonal variations at southern

midlatitudes. The depth of the tropical dip in concentrations of CH3CCl3 will present a test for

the combined effects of intrahemispheric mixing and the disparity between tropical and extra-

tropical levels of OH.

In order to compute the distribution of tropospheric OH presented here we needed to com-

pile distributions for O3 , H2O, NOx , CO, hydrocarbons to the best of our current knowledge.

These distributions, together with those for derived concentrations of selected species, such as

CH2O, members of odd-H and NOx families, and for selected J-values, represent a description,

however imperfect, of the climatological chemical state of the troposphere.

Acknowledgements. We are grateful to the late John Bradshaw and Scott Smyth of the

Georgia Institute of Technology who provided merged data files and gridded data for NO meas-

urements from the GTE and AASE program prior to 1995. Data for PEM-Tropics were obtained

from the GTE data archive (http://www-gte.larc.nasa.gov). We thank Paul Fraser and Paul Nov-

elly and Thayne Thompson (anybody else) for sharing with us unpublished data. We are

indebted to Pauline Midgley for promptly providing essential information on emissions. We ack-

-- --

- 54 -

nowledge important discussions with Daniel Jacob, Hans Schneider, Hanwant Singh, Lyatt Jae-

gle and Martin Schultz. We benefited from expert help in data analysis provided by Amy Mun-

son and Inna Megretskaia. This work was supported by a grant from the National Science Foun-

dation (ATM-9320778).

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- 55 -

References

AFEAS (Alternative Fluorocarbons Environmental Acceptability Study), Production, sales and

atmospheric release of fluorocarbons through 1996, AFEAS, 1333 H Street NW, Washington

DC 20005, USA,1998.

Andronache C., W.L. Chameides, M.O. Rodgers, J. Martinez, P. Zimmerman, J. Greenberg,

Vertical distribution of isoprene in the lower boundary layer of the rural and urban southern

United States, J. Geophys. Res., 99,, 16989-16999, 1994.

Atkinson, R., D.L. Baulch, R.A. Cox, R.F. Hampson, Jr., J.A. Kerr, M. J. Rossi, and J. Troe,

Evaluated kinetic, photochemical and heterogeneous data for atmospheric chemistry, supplement

5. IUPAC subcommittee on gas kinetic data, evaluation for atmospheric chemistry, J. Phys.

Chem Ref. Data, 26, 521-1011, 1997.

Ayers, G. P., and R. W. Gillett, Isoprene emissions from vegetation and hydrocarbon emissions

from bush fires in tropical Australia, J. Atmos. Chem., 7, 177-190, 1988.

Balkanski, Y. J., and D. J. Jacob, Transport of continental air to the subantarctic Indian Ocean,

Tellus, 42, 62-75, 1990.

Balkanski, Y.J., D.J. Jacob, R. Arimoto, and M.A. Kritz, Long-range transport of radon-222

over the North Pacific Ocean: implications for continental influence, J. Atmos. Chem., 14 353-

374, 1992.

Balkanski, Y.J., D.J. Jacob, G.M. Gardner, W.M. Graustein, and K.K. Turekian, Transport and

residence times of continental aerosols inferred from a global 3-dimensional simulation of

210Pb, J. Geophys. Res., 98, 20573-20586, 1993.

Blake D.R., N.J. Blake, T.W. Smith, O.W. Wingenter, F.S. Rowland, Nonmethane hydrocarbon

and halocarbon distributions during Atlantic Stratocumulus Transition Experiment Marine Aero-

-- --

- 56 -

sol and Gas Exchange, June 1992, J. Geophys. Res., 101,, 4501-4514, 1996.

Blake D.R., T-.Y. Chen, T.W. Smith, C.J.L. Wang, O.W. Wingenter, N.J. Blake, F.S. Rowland,

E.W. Mayer, Three-dimensional distribution of nonmenthane hydrocarbons and halocarbons

over the Northwestern Pacific during the 1991 Pacific Exploratory Mission (PEM-WEST A), J.

Geophys. Res., 101,, 1763-1778, 1996.

Blake D.R., D.F. Hurst, T.W. Smith, W.J. Whipple, T-.Y. Chen, N.J. Blake, and F.S. Rowland,

Summertime measurements of selected nonmethane hydrocarbons in the Arctic and Subarctic

during the 1988 Arctic Boundary Layer Expedition (ABLE 3A), J. Geophys. Res., 97,, 16,559-

16588, 1992.

Blake D.R., T.W. Smith, T.Y. Chen, W.J. Whipple, F.S. Rowland, Effects of biomass burning

on summertime nonmethane hydrocarbon concentrations in the Canadian wetlands, J. Geophys.

Res., 99,, 1699-1719, 1994.

Blake N.J., D.R. Blake, T.Y. Chen, J.E. Collins, G.W. Sachse, B.E. Anderson, F.S. Rowland,

Distribution and seasonality of selected hydrocarbons and halocarbons over the western pacific

basin during PEM-WEST A and PEM-WEST B, J. Geophys. Res., 102,, 28315-28331, 1997.

Blake N.J., D.R. Blake, B.C. Sive, T.Y. Chen, F.S. Rowland, J.E. Collins, G.W. Sachse, B.E.

Anderson, Biomass burning emissions and vertical distribution of atmospheric methyl halides

and other reduced carbon gases in the South Atlantic region, J. Geophys. Res., 101,, 24151-

24164, 1996.

Bottenheim, J. W., M. F. Shepherd, C2-C6 hydrocarbon measurements at 4 rural locations across

Canada, Atmos. Environ., 29 647-664, 1995.

Bradshaw, J., S. B. Smyth, S. C. Liu, R. Newell, D. D. Davis, and S. T. Sandholm, Observed dis-

tributions of nitrogen oxides in the remote free troposphere from NASA Global Tropospheric

Chemistry Experiment, Reviews of Geophysics, in press, 1998.

-- --

- 57 -

Brenninkmeijer, C.A.M., M.R. Manning, D.C. Lowe, G. Wallace, R.J. Sparks, and A. Volz-

Thomas, Interhemispheric asymmetry in OH abundance inferred from measurements of atmos-

pheric 14CO, Nature, 356, 50-52, 1992.

Brenninkmeijer, C.A.M., Measurement of the abundance of 14CO in the atmosphere and the

13C/12C and 18O/16O ratio of atmospheric CO, with applications in New Zealand and

Antarctica, J. Geophys. Res., 98, 10,595-10,614, 1993.

Brune, W.H., I.C. Faloona, D. Tan, A.J. Weinheimer, T. Campos, B.A. Ridley, S.A. Vay, J.E.

Collins, G.W. Sachse, L. Jaegle, D.J. Jacob, Airborne in-situ OH and HO2 observations in the

cloud-free troposphere and lower stratosphere during SUCCESS, Geophys. Res. Lett., 25,,

1,701-1,704, 1998.

Butler, J.H., J. W. Elkins, T. M. Thompson, B. D. Hall, T. H. Swanson, and V. Koropalov, Oce-

anic consumption of CH3CCl3: Implications for tropospheric OH, J. Geophys. Res., 96, 22347-

22355, 1991.

Carroll, M.A., and A. M. Thompson, NOx in the non-urban troposphere, in Problems and Pro-

gress in Atmospheric Chemistry, ed. J. Barker, Adv. in Phys. Chem., Vol. 3, World Scientific

Pub. Co., London, 1995.

Chameides, W. L. and A. Tan, The two-dimensional diagnostic model for tropospheric OH: an

uncertainty analysis, J. Geophys. Res., 86, 5209-5223, 1981.

Chin, M., and D.J. Jacob, Anthropogenic and natural contributions to atmospheric sulfate: a glo-

bal model analysis, J. Geophys. Res., 1 01, 18,691-18,700, 1996.

Chin, M., D.J. Jacob, G.M. Gardner, M. Foreman-Fowler, P.A. Spiro, and D. L. Savoie, A global

three-dimensional model of tropospheric sulfate, J. Geophys. Res. 101, 18,667-18,690, 1996.

Crutzen, P. J., and L. T. Gidel, A two-dimensional photochemical model of the atmosphere 2:

-- --

- 58 -

the tropospheric budgets of the anthropogenic chlorocarbons CO, CH4 , CH3Cl and the effect of

various NOx sources on tropospheric ozone, J. Geophys. Res., 88, 6641-6661, 1983.

Crutzen, P. J., and J. Fishman, Average concentrations of OH in the troposphere and the budgets

of CH4 , CO, H2 , and CH3CCl3 , Geophys. Res. Lett., 4,, 321-324, 1977.

Crutzen, P.J. and P. H. Zimmermann, The changing photochemistry of the troposphere, Tellus,

43AB, 136-151, 1991.

Cunnold, D., P. Fraser, R. Weiss, R. Prinn, P. Simmonds, F. Alyea, and A. Crawford. Global

trends and annual releases of CCl3F and CCl2F2 estimated from ALE/GAGE and other meas-

urements from July 1978 to June 1991. J. Geophys. Res. 99, 1107-1126, 1994.

DeMore, W. B., S. P. Sander, D. M. Golden, R. F. Hampson, M. J. Kurylo, C. J. Howard, A. R.

Ravishankara, C. E. Kolb, and M. J. Molina, Chemical kinetics and photochemical data for use

in stratospheric modeling, Evaluation No. 9, JPL Publication 94−1, Jet Propulsion Lab.,

Pasadena, CA, 1994.

DeMore, W.B., S.P. Sander, D.M. Golden, R.F. Hampson, M.J. Kurylo, C.J. Howard, A.R. Rav-

ishankara, C.E. Kolb, and M.J. Molina, Chemical kinetics and photochemical data for use in

stratospheric modeling, JPL publication 97−4, Pasadena, California, 1997.

Dlugokencky, E. J., K. A. Masarie, P. M. Lang, P. P. Tans, Continuing decline in the growth rate

of the atmospheric methane burden, Nature, 393, 447-450, 1998.

Dlugokencky, E. J., L. P. Steele, P. M. Lang, K. A. Masarie, Atmospheric methane at Mauna

Loa and Barrow observatories - presentation and analysis of in situ measurements, J. Geophys.

Res., 100, 23103-23113, 1995.

Dlugokencky, E. J., L. P. Steele, P. M. Lang, K. A. Masarie, The growth rate and distribution of

atmospheric methane, J. Geophys. Res., 99, 17021-17043, 1994.

-- --

- 59 -

Donahue, N. M., R. G. Prinn, Insitu nonmethane hydrocarbon measurements on SAGA-3, J.

Geophys. Res., 98, 16915-16932, 1993.

Drummond, J.W., D.H. Ehhalt, and A. Volz, Measurements of nitric oxide between 0-12 km alti-

tude and 67° N to 60° S latitude obtained during STRATOZ III, J. Geophys. Res., 93, 15,831-

15,849, 1988.

Ehhalt, D.H., H.P. Dorn, D. Poppe, The chemistry of the hydroxyl radical in the troposphere, P

ROY SOC EDINB B 97: 17-34 1990.

Ehhalt, D.H., and J. Drummond, NOx sources and the tropospheric distribution of NOx during

STRATOZ III, Proceedings of the NATO Advanced Research Workshop on Regional and Global

Ozone and its Environmental Consequences, ed. by I.S.A. Isaksen, NATO ASI Series C, 227,

pp217-237, Reidel, Dordrecht, Holland, 1988.

Eisele F. L., G. H. Mount, D. Tanner, A. Jefferson, R. Shetter, J. W. Harder, and E. J. Williams,

Understanding the production and interconversion of the hydroxyl radical during the tropos-

pheric OH photochemistry experiment, J. Geophys. Res., 102, 6457-6465, 1997.

Eluszkiewicz, J., D. Crisp, R. Zurek, L. Elson, L. Froidevaux, J. Waters, R.G. Grainger, A. Lam-

bert, R. Harwood, and G. Peckham, Residual circulation in the stratosphere and lower meso-

sphere as diagnosed from Microwave Limb Sounder data, J. Atmos. Sci., 53, 217-240, 1996.

Emmons, L. K., M. A. Carroll, et al., Climatologies of NOx and NOy: A comparison of data and

models, Atmos. Environ. 31, 1851-1904, 1997.

Fan, S.-M., D. J. Jacob, D. L. Mauzerall, J. D. Bradshaw, S. T. Sandholm, D. R. Blake, H. B.

Singh, R. W. Talbot, G. L. Gregory, and G. W. Sachse, Origin of tropospheric NOx over subarc-

tic eastern Canada in summer, J. Geophys. Res., 99, 16867-16877, 1994.

Fisher, D., NASA Report on Concentrations, Lifetimes and Trends of CFC’s, Halons and Related

-- --

- 60 -

Species, chap 2, 1994.

Fishman J., and V.G. Brackett, The climatological distribution of tropospheric ozone derived

from satellite measure ments using version 7 Total Ozone Mapping Spectrometer and Stratos-

pheric Aerosol and Gas experiment data sets, J. Geophys. Res. 102, 19,275-19,278, 1997.

Fishman, J., S. Solomon, and P.J. Crutzen, Observational and theoretical evidence in support of

a significant in-situ photochemical source of tropospheric ozone, Tellus, 31, 432-446, 1979.

Fishman, J., C.E. Watson, J.C. Larsen, and J.A. Logan, The distribution of tropospheric ozone

obtained from satellite data, J. Geophys. Res., 95, 3599-3617, 1990.

Fraser, P. J., P. Hyson, R. A. Rasmussen, A. J. Crawford, and M. A. K. Khalil, Methane, carbon

monoxide and methylchloroform in the southern hemisphere, J. Atmos. Chem., 4, 3-42, 1986.

Goldan, P. D., W. C. Kuster, F. C. Fehsenfeld, and S. A. Montzka, Hydrocarbon measurements

in the southeastern United States: The rural oxidants in the southern environment program 1990,

J. Geophys. Res., 100, 25945-25963, 1995.

Goldstein, A.H., S. C. Wofsy, and C. M. Spivakovsky, Seasonal variations of nonmethane

hydrocarbons in rural New England: constraints on OH concentrations in northern mid-latitudes,

J. Geophys. Res. 100, 21,023-21,033, 1995a.

Goldstein, A. H., B. C. Daube, J. W. Munger, S. C. Wofsy, Automated in-situ monitoring of

atmospheric non-methane hydrocarbon concentrations and gradients, J. Atm. Chem., 21, 43-59,

1995b.

Guenther, A., W. Baugh, K. Davis, G. Hampton, P. Harley, L. Klinger, L. Vierling, P. Zimmer-

man, E. Allwine, S. Dilts, B. Lamb, H. Westberg, D. Baldocchi, C. Geron, and T. Pierce,

Isoprene fluxes measured by enclosure, relaxed eddy accumulation, surface layer gradient,

mixed layer gradient, and mixed layer mass balance techniques, J. Geophys. Res. 101, 18555-

-- --

- 61 -

18567, 1996.

Hansen, J., G. Russell, D. Rind, P. Stone, A. Lacis, S. Lebedeff, R. Ruedy, and L. Travis,

Efficient three-dimensional global models for climate studies: models I and II, Monthly Weather

Rev., 111, 609-662, 1983.

Hao, W.M., M.H. Liu, and P.J. Crutzen, Estimates of annual and regional releases of CO2 and

other trace gases to the atmosphere from fires in the tropics, based on the FAO statistics for the

period 1975-1980, in Proceedings of Third International Symposium on Fire Ecology, Freiburg

University, Federal Republic of Germany 16-10, May 1989, Springer-Verlag, Berlin, 1990.

Helmig, D., B. Balsey, K. Davis, L. R. Kuck, M. Jensen, J. Bognar, T. Smith Jr., R. V. Arrieta,

R. Rodriguez, and J. W. Birks, Vertical profiling and determination of landscape fluxes of

biogenic nonmethane hydrocarbons within the planetary boundary layer in the Peruvian Ama-

zon, J. Geophys. Res., 103, 25519-25532, 1998.

Holton, J. R., P. H. Haynes, M. E. McIntyre, A. R. Douglas, R. B. Rood, and L. Pfister,

Stratosphere-troposphere exchange, Rev. Geophys. 33, 403-439, 1995.

Horowitz, L.W., J. Liang, G.M. Gardner, and D.J. Jacob, Export of reactive nitrogen from North

America during summertime: Sensitivity to hydrocarbon chemistry, J. Geophys. Res., 103,

13,451-13,476, 1998.

S. Houweling, F. Dentener, J. Lelieveld, The impact of nonmethane hydrocarbon compounds on

tropospheric photochemistry, J. Geophys. Res., 103,, 10673-10696, 1998.

Jackman, C.H., A. R. Douglass, K. F. Brueske and S. A. Klein, The influence of dynamics on

two-dimensional model results: simulation of 14C and stratospheric aircraft NOx injections, J.

Geophys. Res., 96,, 22,559-22572, 1991.

Jacob, D. J., and M. J. Prather, Radon-222 as a test of convective transport in a general circula-

-- --

- 62 -

tion model, Tellus, 42,118-134,1990.

Jacob, D. J., M. J. Prather, S. C. Wofsy, and M. B. McElroy, Atmospheric distribution of 85Kr

simulated with a general circulation model, J. Geophys. Res., 92, 6614-6626, 1987.

Jacob, D. J., and S. C. Wofsy. Photochemistry of biogenic emissions over the Amazon forest, J.

Geophys. Res., 93, 1477-1486, 1988a.

Jacob, D.J., and S.C. Wofsy, Photochemical production of carboxylic acids in a remote continen-

tal atmosphere, in Acid Deposition at High Elevation Sites, ed. by M.H. Unsworth and D.

Fowler, Kluwer Academic Publishers, Dordrecht, 1988b.

Jaegle L., D. J. Jacob, P. O. Wennberg, C. M. Spivakovsky, T.F. Hanisco, E. J. Lanzendorf, E. J.

Hintsa, D. W. Fahey, E. R. Keim, M. H. Proffitt, E. L. Atlas, F. Flocke, S. Schauffler, C. T.

Mcelroy, C. Midwinter, L. Pfister, J. C. Wilson, Observed OH and HO2 in the upper troposphere

suggest a major source from convective injection of peroxides, Geophys. Res. Lett., 24, 3181-

3184, 1997.

Jobson B. T., D. D. Parrish, P. Goldan, W. Kuster, F. C. Fehsenfeld, D. R. Blake, N. J. Blake, H.

Niki, Spatial and temporal variability of nonmethane hydrocarbon mixing ratios and their rela-

tion to photochemical lifetime, J. Geophys. Res.,103, 13557-13567, 1998.

Jobson B. T., Z. Wu, H. Niki, L. A. Barrie, Seasonal trends of isoprene, C-2-C-5 alkanes, and

acetylene at a remote boreal site in Canada, J. Geophys. Res., 99,1589-1599, 1994.

Johnston, H., Evaluation of excess carbon 14 and strontium 90 data for suitability to test two-

dimensional stratospheric models, J. Geophys. Res., 94, 18,485-18, 18,493, 1989.

Kanakidou M, F. J. Dentener, P.J. Crutzen, A global three-dimensional study of the fate of

HCFCs and HFC-134A in the troposphere, J. Geophys. Res., 100,18,781-18,801, 1995.

Khalil, M. A. K., and R. A. Rasmussen, The atmospheric lifetime of methylchloroform

-- --

- 63 -

(CH3CCl3), Tellus, 36B, 317-332, 1984.

Koch D. M., D. J. Jacob, W. C. Graustein, Vertical transport of tropospheric aerosols as indi-

cated by BE-7 and PB-210 in a chemical tracer model, J. Geophys. Res.,101, 18651-18666,

1996.

Koppmann R., R. Bauer, F.R. Johnen, C. Plass, and J. Rudolph, The distribution of light non-

methane hydrocarbons over the mid-Atlantic: results of the Polarstern cruise ANT VII/1, J.

Atmos. Chem., 15,215-15,234, 1992.

Koppmann R., F. J. Johnen, C. Plass-Dulmer and J. Rudolph, Distribution of methylchloride,

dichloromethane, trichloroethene and tetrachloroethene over the north and south Atlantic, J.

Geophys. Res. 98, 20,517-20,526, 1993.

Kondo, Y., T. Kitada, M. Koike, S. Kawakami, and Y. Makino, Nitric oxide and ozone in the

free troposphere over the western Pacific ocean, J. Geophys. Res. 98, 20,527-20,535, 1993.

Krol M., P. J. van Leeuwen, and J. Lelieveld, Global OH trend inferred from methylchloroform

measurements, J. Geophys. Res., 103, 10,697-10,711, 1998.

Levy, H., II, Normal atmosphere: large radical and formaldehyde concentrations predicted, Sci-

ence, 173, 141-143, 1971.

Lingenfelter, R.E., Production of carbon 14 by cosmic-ray neutrons, Rev. Geophys., 1, 35-55,

1963.

Lindskog, A., J. Moldanova, The influence of the origin, season and time of the day on the distri-

bution of individual NMHC measured at Rorvik, Sweden, Atmos. Environ., 28, 2383-2398,

1994.

Liu, W. T., and W. Tang, Precipitable water and surface humidity over global oceans from spe-

cial sensor microwave imager and European Center for Medium Range Weather forcasts, J.

-- --

- 64 -

Geophys. Res.,97, 2251-2264, 1992.

Logan, J. A., An analysis of ozonesonde data for the troposphere: recommendations for testing

3-D models, and development of a gridded climatology for tropospheric ozone, J. Geophys.

Res., in press, 1999.

Logan, J. A., M. J. Prather, S. C. Wofsy, and M. B. McElroy, Tropospheric chemistry: a global

perspective, J. Geophys. Res., 86, 7210-7254, 1981.

Lovelock, J. E., Methyl chloroform in the troposphere as an indicator of OH radical abundance,

Nature, 267, 32-33, 1977.

Mak, J.E., C.A.M. Brenninkmeijer, and M.R. Manning, Evidence for a missing carbon monoxide

sink based on tropospheric measurements of 14CO, Geophys. Res. Lett., 19, 1467-1470, 1992.

Mak, J.E., C.A.M. Brenninkmeijer, and J. Tamaresis, Atmospheric 14CO observations and their

use for estimating carbon monoxide removal rates, J. Geophys. Res. 99, 22,915-22,922, 1994.

Makide, Y., F. S. Rowland, Tropospheric concentrations of methylchloroform, CH3CCl3 , in

January 1978 and estimates of atmospheric residence times for hydrohalocarbons, Proceedings,

National Academy of Sciences (U.S.A.) 78, 5933-5973, 1981.

Manning, M. R., C. A. M. Brenninkmeijer, W. Allan, Atmospheric carbon monoxide budget of

the southern hemisphere - implications of C-13/C-12 measurements, J. Geophys. Res. 102,

10673-10682, 1997.

Mather, J. H., P. S. Stevens, W.H. Brune, OH and HO2 measurements using laser-induced

fluorescence, J. Geophys. Res., 102, 6427-6436, 1997.

McConnell, J. C., M. B. McElroy and S. C. Wofsy, Natural sources of atmospheric CO, Nature,

233, 187-188, 1971.

-- --

- 65 -

McCormick, M.P., E.W. Chiou, L.R. McMaster, W. P. Chu, J.C. Larsen, D. Rind, and S. Olt-

mans, Annual variation of water vapor in the stratosphere and upper troposphere observed by the

Stratospheric Aerosol and Gas Experiment, J. Geophys. Res., 98, 4867-4875, 1993.

McCulloch, A., P.M. Midgley and D.A. Fisher, Distribution of emissions of chlorofluorocarbons

(CFCs) 11, 12, 113, 114 and 115 among reporting and non-reporting countries, Atmospheric

Environment 28, 2567-2582, 1994.

McCulloch A., P.M. Midgley, The production and global distribution of emissions of tri-

chloroethene, tetrachloroethene, and dichloromethane over the period 1988-1992, Atmos.

Environ., 30, 601-608, 1996.

Mckeen, S. A., G. Mount, F. Eisele, E. Williams, J. Harder, P. Goldan, W. Kuster, S. C. Liu, K.

Baumann, D. Tanner, A. Fried, S. Sewell, C. Cantrell, and R. Shetter, Photochemical modeling

of hydroxyl and its relationship to other species during the tropospheric OH photochemistry

experiment, J. Geophys. Res., 102, 6467-6493, 1997.

Michelsen, H. A., R. J. Salawitch, P. O. Wennberg, and J. G. Anderson, Production of O(1D)

from photolysis of O3 , Geophys. Res. Lett., 21, 2227-2230, 1994.

Midgley, P.M., The production and release to the atmosphere of 1,1,1-trichloroethane (methyl

chloroform), Atmos. Environ., 23 2663-2665, 1989.

Midgley, P.M. and A. McCulloch, The production and global distribution of emissions to the

atmosphere of 1,1,1-trichloroethane (methyl chloroform), Atmos. Environ., 29, 1601 - 1608,

1995.

Midgley, P.M. and A. McCulloch, Estimated national releases to the atmosphere of

chlorodifluoromethane (HCFC-22) during 1990, Atmos. Environ., 31, 809-811, 1997.

Miller, B.R., J. Huang, R.F. Weiss, R.G. Prinn, P.J. Fraser, Atmospheric trend and lifetime of

-- --

- 66 -

chlorodifluoromethane (HCFC-22) and the global tropospheric OH concentration, J. Geophys.

Res., 103, 13,237-13,248, 1998.

Montzka, S. A., J. H. Butler, R. C. Myers, T. M. Thompson, T. H. Swanson, A. D. Clarke, L. T.

Lock, J. W. Elkins, Decline in the tropospheric abundance of halogen from halocarbons: Impli-

cations for stratospheric ozone depletion, Science, 272, 1318-1322, 1996.

Montzka, S. A., R. C. Myers, J. H. Butler, J. W. Elkins, and S. O. Cummings, Global tropos-

pheric distribution and calibration scale of HCFC-22, Geophys Res. Lett., 20, 703-706, 1993.

Mount G. H., J. W. Brault, P. V. Johnston, E. Marovich, R. O. Jakoubek, C. J. Volpe, J.

Harder, and J. Olson, Measurement of tropospheric OH by long-path laser absorption at Fritz

Peak observatory, Colorado, during the OH photochemistry experiment, fall 1993, J. Geophys.

Res., 102, 6393-6413, 1997.

Muller, J.-F., and G. Brasseur, IMAGES: A three-dimensionsal chemical transport model of the

global troposphere, J. Geophys. Res., 100, 16,445-16490, 1995.

Munger, J. W., S. M. Fan, P. S. Bakwin, M. L. Goulden, A. H. Goldstein, A. S. Colman, S. C.

Wofsy, Regional budgets for nitrogen oxides from continental sources - variations of rates for

oxidation and deposition with season and distance from source regions, J. Geophys. Res., 103,

8355-8368, 1998.

Novelli, P.C., L.P. Steele and P. Tans, Mixing ratios of carbon monoxide in the troposphere, J.

Geophys. Res., 97, 20,731- 20,750, 1992.

Novelli, P. C., K. A. Masarie, P. P. Tans, and P. M. Lang, Recent changes in atmospheric carbon

monoxide, Science, 263, 1587-1590, 1994.

O’Brien, K., Secular variations in the productions of cosmogenic isotopes in the Earth’s atmo-

sphere, J. Geophys. Res., 84, 423-432, 1979.

-- --

- 67 -

Oltmans, S. J., and W. D. Komhyr, Surface ozone distributions and variations for 1973-1984

measurements at the NOAA Geophysical Monitoring for Climatic Change baseline observa-

tories, J. Geophys. Res., 91, 5229-5236, 1986.

Oort, A. H., Global atmospheric circulation statistics, 1958-1973. NOAA Professional Paper 14.

U. S. Dept. of Commerce, Rockville, MD, 1983.

Paulson, S.E. and J.H. Seinfeld, Development and evaluation of a photooxidation mechanism for

isoprene, J. Geophys. Res., 97, 20,703-20,715, 1992.

Peixoto, J. P., and A. H. Oort, Physics of climate, American Institute of physics, New York,

1992.

Penkett, S.A., and K.A. Brice, The spring maximum in photo-oxidants in the Northen Hemi-

sphere troposphere, Nature, 319, 655-657, 1986.

Penkett, S.A., N.J. Blake, P. Lightman, A.R.W. Marsh, P. Anwyl, and G. Butcher, The seasonal

variation of non-methane hydrocarbons in the free troposphere over the North Atlantic Ocean:

possible evidence for extensive reaction of hydrocarbons with the nitrate radical, J. Geophys.

Res., 98, 2865-2886, 1993.

Prather, M. J., Solution of the inhomogeneous Rayleigh scattering atmosphere, The Astrophysi-

cal Journal, 192, 787-792, 1974

Prather, M. J., D. J. Jacob, A persistent imbalance in HOX and NOX photochemistry of the

upper troposphere driven by deep tropical convection, Geophys Res. Lett., 24, 3189-3192, 1997.

Prather, M. J., M. B. McElroy, S. C. Wofsy, G. Russell, and D. Rind, Chemistry of the global

troposphere: fluorocarbons as tracers of air motion, J. Geophys. Res., 92, 6579-6613, 1987.

Prather, M. J., and E. E. Remsberg, The atmospheric effects of aircraft: Report of the 1992

Models and Measurements Workshop, NASA Reference Publication 1292, 1993.

-- --

- 68 -

Prather, M. J., and C. M. Spivakovsky, Tropospheric OH and the lifetimes of

hydrochlorofluorocarbons, J. Geophys. Res., 95, 18,723-18,729, 1990.

Prinn, R. G., R. A. Rasmussen, P. G. Simmonds, F. N. Alyea, D. M. Cunnold, B. C. Lane, C. A.

Cardelino, and A. J. Crawford, The atmospheric lifetime experiment, 5: results for CH3CCl3

based on three years of data, J. Geophys. Res., 88, 8415-8426, 1983.

Prinn, R., D. Cunnold, R. Rasmussen, P. Simmonds, F. Alyea, A. Crawford, P. Fraser, and R.

Rosen, Atmospheric trends in methylchloroform and the global average for the hydroxyl radical,

Science, 238, 945-950, 1987.

Prinn, R. G., D. M. Cunnold, P. G. Simmonds, F. N. Alyea, R. Boldi, D. Gutzler, D. Hartley, R.

Rosen, and R.A. Rasmussen, Global average concentration and trend for hydroxyl radicals

deduced from ALE/GAGE trichloroethane (methyl chloroform) data for 1978-90, J. Geophys.

Res., 97, 2445-2461, 1992.

Prinn R. G., R. F. Weiss, B. R. Miller, J. Huang, F. N. Alyea, D. M. Cunnold, P. J. Fraser, D. E.

Hartley, and P. G. Simmonds, Atmospheric trends and lifetime of CH3CCl3 and global OH con-

centrations, Science 269, 187-192, 1995.

Rasmussen, R. A., and M. A. K. Khalil, Isoprene over the Amazon basin, J. Geophys. Res., 93,

1417-1421, 1988.

Ravishankara A. R., D. L. Albritton, Methyl chloroform and the atmosphere, Science, 269,183-

184, 1995.

Ridley, B.A., J. G. Walega, J. E. Dye, and F. E. Grahek, Distributions of NO, NOx , NOy , and O3

to 12 km altitude during the summer monsoon season over New Mexico, J. Geophys. Res.,99,

25,519-25,534, 1994.

Rind, D., and J. Lerner, Use of on-line tracers as a diagnostic tool in general circulation model

-- --

- 69 -

development, J. Geophys.Res., 101, 12,667-12,683, 1996.

Roelofs, G. J., and J. Lilieveld, Distribution and budget of tropospheric ozone calculated iwth a

chemistry general circulation model, J. Geophys. Res., 100, 20,983-20998, 1995.

Rohrer, F., D. Bruning, and D. H. Ehhalt, Tropospheric Mixing ratios of NO obtained during

TROPOZ II in the latitude region 67°N-56°S, J. Geophys. Res., 102, 25,429-25,449, 1997.

Rosenlof K. H., and J. R. Holton, Estimates of the stratospheric residual circulation using the

downward control principle, J. Geophys. Res., 98, 10,456-10,479, 1993.

Rossow, W.B., and R. A. Schiffer, ISCCP cloud data products, Bulletin American Meteorologi-

cal Society, 72, 1-20, 1991.

Rudolph J., Two-dimensional distribution of light hydrocarbons: results from the STRATOZ III

experiment, J. Geophys. Res., 93, 8367-8377, 1988.

Rudolph J., The tropospheric distribution and budget of ethane, J. Geophys. Res., 100, 11,369-

11381, 1995.

Rudolph J., and F.J. Johnen, Measurements of light hydrocarbons over the Atlantic in regions of

low biological activity, J. Geophys. Res., 95, 20,583-20,591, 1990.

Rudolph, J., A., Khedim, and B. Bonsag, Light hydrocarbons in the tropospheric boundary layer

over tropical Africa, J. Geophys. Res., 97, 6181-6186, 1992a.

Rudolph, J., A., Khedim, T. Clarkson, and D. Wagenbach, Long term measurements of light

alkanes and acetylene in the Antarctic troposphere, Tellus, 44(B), 252-261, 1992b

Rudolph, J., A., Khedim, R. Koppmann, B. Bonsang, Field Study of the Emissions of Methyl

Chloride and Other Halocarbons from Biomass Burning in Western Africa, J. Atmos. Chem., 22,

67-80, 1995.

-- --

- 70 -

Rudolph J., R. Koppmann, C. Plassdulmer, The budgets of ethane and tetrachloroethene - is

there evidence for an impact of reactions with chlorine atoms in the troposphere, Atmos.

Environ., 30, 1887-1894, 1996.

Schneider H. R., D. B. Jones, G.-Y. Shi, and M. B. McElroy, Analysis of residual mean transport

in the stratosphere. Part I: Model description and comparison with satellite data, J. Geophys.

Res.,, submitted, 1998.

Shia, R.-L., Y. L. Yung, M. Allen, R. Zurek and D. Crisp, Sensitivity study of advection and dif-

fusion coefficients in a two-dimensional stratospheric model using excess carbon 14 data, J.

Geophys, Res., 94, 18,467-18484, 1989.

Singh, H. B., Atmospheric halocarbons: evidence in favor of reduced hydroxyl radical concen-

trations in the troposphere, Geophys. Res. Lett., 4, 241-244, 1977a.

Singh H. B., Preliminary estimation of average tropospheric HO concentrations in the northern

and southern hemispheres, Geophys. Res. Lett., 4, 453-456, 1977b.

Singh H.B., D. Herlth, R. Kolyer, R. Chatfield, W. Viezee, L.J. Salas, Y. Chen, J.D. Bradshaw,

S.T. Sandholm, R. Talbot, G.L. Gregory, B. Anderson, G.W. Sachse, E. Browell, A.S.

Bachmeier, D.R. Blake, B. Heikes, D. Jacob, H.E. Fuelberg, Impact of biomass burning emis-

sions on the composition of the south atlantic troposphere - reactive nitrogen and ozone, J. Geo-

phys. Res., 101, 24203-24219, 1996.

Singh H.B., D. Herlth, R. Kolyer, L. Salas, J.D. Bradshaw, S.T. Sandholm, D.D. Davis, J. Craw-

ford, Y. Kondo, M. Koike, R. Talbot, G. L. Gregory, G.W. Sachse, E. Browell, D.R. Blake, F.S.

Rowland, R. Newell, J. Merrill, B. Heikes, S.C. Liu, P.J. Crutzen, M. Kanakidou, Reactive nitro-

gen and ozone over the western pacific - distribution, partitioning, and sources, J. Geophys. Res.,

101, 1793-1808, 1996.

Singh H. B., D. Herlth, D. Ohara, K. Zahnle, J.D. Bradshaw, S.T. Sandholm, R. Talbot, G.L.

-- --

- 71 -

Gregory, G.W. Sachse, D.R. Blake, S.C. Wofsy, Summertime distribution of pan and other reac-

tive nitrogen species in the northern high-latitude atmosphere of eastern Canada, J. Geophys.

Res., 99, 1821-1835, 1994.

Singh, H. B., M. Kanakidou, P. J. Crutzen, D. J. Jacob, High concentrations and photochemical

fate of oxygenated hydrocarbons in the global troposphere, Nature, 378, 50-54, 1995.

Singh H.B., D. Ohara, D. Herlth, W. Sachse, D.R. Blake, J.D. Bradshaw, M. Kanakidou, P.J.

Crutzen, Acetone in the atmosphere - distribution, sources, and sinks, J. Geophys. Res., 99,

1805-1819, 1994.

Singh, H.B., A.N. Thakur, Y.E. Chen, and M. Kanakidou, Tetrachloroethylene as an indicator of

low Cl atom concentrations in the troposphere, Geophys. Res. Lett., 23,, 1529-1532, 1996.

Singh, H.B., and 10 others, Summertime distribution of PAN and other reactive nitrogen species

in the northern high-latitude atmosphere of eastern Canada, J. Geophys. Res., 99, 1821-1836,

1994.

Singh, H. B. and 21 others, Reactive nitrogen and ozone over the western Pacific - distribution,

partitioning, and sources , J. Geophys, Res., 101 1793-1808, 1996.

Singh, H.B., W. Viezee and L.J. Salas, Measurements od selected C2-C5 hydrocarbons in the

troposphere: Latitudinal, vertical and temporal variations, J. Geophys. Res. 93, 15,861-15,878,

1988.

Sobolev, V. V., Light scattering in planetary atmospheres, Oxford, New York, Pergamon Press

[1975]

Soden, B. J., F. P Bretherton, Evaluation of water vapor distribution in general circulation

models using satellite observations, J. Geophys, Res., 99, 1187-1210, 1994.

Spivakovsky, C.M., and Y.J. Balkanski, Tropospheric OH: constraints imposed by observations

-- --

- 72 -

of 14CO and CH3CCl3, Report of the WMO-sponsored meeting of carbon monoxide (CO)

experts, P.C. Novelli and R.M. Rosson (eds.), Global Atmospheric Watch, World Meteorologi-

cal Organization, 1994.

Spivakovsky, C.M., R. Yevich, J.A. Logan, S.C. Wofsy, M.B. McElroy, and M. J. Prather, Tro-

pospheric OH in a three dimensional chemical tracer model: an assessment based on observa-

tions of CH3CCl3 , J. Geophys. Res., 95, 18,441-18,471, 1990.

Spivakovsky, C.M., Reply, J. Geophys. Res. 96, 17395-17398, 1991.

Talukdar, R. K., J. B. Burkholder, M. Hunter, M. K. Gilles, J. M. Roberts, and A. R. Ravishan-

kara, Atmospheric fate of several alkyl nitrates: 2. UV absorption cross-sections and photodisso-

ciation quantum yields, Journal of the Chemical Society-Faraday Transactions 93, 2797-2805,

1997a.

Talukdar, R. K., S. C. Herndon, J. B. Burkholder, J. M. Roberts, and A. R. Ravishankara,

Atmospheric fate of several alkyl nitrates: 1. rate coefficients of the reactions alkyl nitrates with

isotopically labelled hydroxyl radicals, Journal of the Chemical Society-Faraday Transactions

93, 2787-2796, 1997b.

Talukdar, R. K., C. A. Longfellow, M. K. Gilles, and A. R. Ravishankara, Quantum yields of

O(D-1) in the photolysis of ozone between 289 and 329 nm as a function of temperature, Geo-

phys. Res. Lett., 25, 143-146, 1998.

Talukdar, R. K., A. Mellouki, A.-M. Schmoltner, T. Watson, S. Montzka, and A. R. Ravishan-

kara, Kinetics of the OH reaction with methyl chloroform and its atmospheric implications, Sci-

ence, 257, 227-230, 1992.

Tanner D. J., A. Jefferson, F. L. Eisele, Selected ion chemical ionization mass spectrometric

measurement of OH, J. Geophys. Res., 102, 6415-6425, 1997.

-- --

- 73 -

Thompson, A. M., The oxidizing capacity of the earth’s atmosphere: Probable past and future

changes, Science, 256, 1157-1165, 1992.

Thompson, A. M. and R. J. Cicerone, Atmospheric CH4 , CO and OH from 1860-1985. Nature,

321, 148-150, 1986a.

Thompson, A. M. and R. J. Cicerone, Possible perturbations to atmospheric CO, CH4 , and OH,

J. Geophys. Res., 91, 10,853-10,864, 1986b.

Thompson, A. M., M. A. Huntly and R. W. Stewart, Perturbations to tropospheric oxidants,

1985-2035, 1. Calculations of ozone and OH in chemically coherent regions, J. Geophys. Res.,

95, 9829-9844, 1990.

Thompson, A. M., and R. W. Stewart, Effect of chemical kinetics uncertainties on calculated

constituents in a tropospheric photochemical model, fIJ. Geophys. Res., 96, 13,089-13,108,

1991.

Thompson, A. M., R. W. Stewart, M. A. Owens and J. A. Herwehe, Sensitivity of tropospheric

oxidants to chemical and climate change, Atmos. Environ, 23,519-532, 1989.

Trenberth, K. E., Global analyses from ECMWF and atlas of 1000 to 10 mb circulation statistics,

NCAR/TN-373+STR, Natl. Center for Atmos. Res., Boulder, CO, 1992.

Torres, A.L., and H. Buchan, Tropospheric nitric oxide measurements over the Amazon Basin, J.

Geophys. Res., 93, 1396-1406, 1988.

Torres, A.L., and A. M. Thompson, Nitric oxide in the equatorial Pacific boundary layer: SAGA

3 measurements, J. Geophys. Res., 98, 16,949-16,954, 1993.

Volk C. M., J. W. Elkins, D. W. Fahey, G. S. Dutton, J. M. Gilligan, M. Loewenstein, J. R.

Podolske, K. R. Chan, and M. R. Gunson, Evaluation of source gas lifetimes from stratospheric

observations, J. Geophys. Res. 102, 25,543-25,564.

-- --

- 74 -

Volz, A., D. H. Ehhalt, and R. G. Derwent, Seasonal and latitudinal variation of 14CO and the

tropospheric concentration of OH radicals, J. Geophys. Res., 86, 5163-5171, 1981.

Wang, Y. H., D. J. Jacob, and J. A. Logan, Global simulation of Tropospheric O3-NOx-

Hydrocarbon chemistry, 1. Model formulation, J. Geophys. Res. 103, 10,713-10,726, 1998a.

Wang, Y. H., J. A. Logan, and D. J. Jacob, Global simulation of Tropospheric O3-NOx-

Hydrocarbon chemistry, 2. Model evaluation, J. Geophys. Res. 103, 10,727-10,756, 1998b.

Wang, Y. H., D. J. Jacob, and J. A. Logan, Global simulation of Tropospheric O3-NOx-

Hydrocarbon chemistry, 3. Origin of tropospheric ozone and effects of non-methane hydrocar-

bons, J. Geophys. Res. 103, 10,757-10,768, 1998c.

Warneck, P., On the role of OH and HO2 radicals in the troposphere, Tellus, 26, 39-46, 1974.

Weinstock, B., and H. Niki, Carbon monoxide balance in nature, Science, 176, 290-292, 1972.

Wennberg, P. O., R. C. Cohen, R. M. Stimpfle, J. P. Koplow, J. G. Anderson, R. J. Salawitch, D.

W. Fahey, E. L. Woodbridge, E. R. Keim, R. S. Gao, C. R. Webster, R. D. May, D. W. Toohey,

L. M. Avallone, M. H. Proffitt, M. Loewenstein, J. R. Podolske, K. R. Chan, S. C. Wofsy,

Removal of stratospheric O-3 by radicals - in situ measurements of OH, HO2, NO, NO2, ClO,

and BrO, Science, 266, 398 -404, 1994.

Wennberg, P. O., T. F. Hanisco, L. Jaegle, D. J. Jacob, E. J. Hintsa, E.J Lanzendo, J. G. Ander-

son, R. S. Gao, E. R. Keim ER, S. G. Donnelly, L. A. Delnegro, D. W. Fahey, S. A. Mckeen, R.

J. Salawitch, C. R. Webster, R. D. May, R. L. Herman, M. H. Proffitt, J. J. Margitan, E. L. Atlas,

S. M. Schauffler, F. Flocke, C. T. Mcelroy, T. P. Bui, Hydrogen radicals, nitrogen radicals, and

the production of O-3 in the upper troposphere, Science 279, 49-53, 1998.

Weiss, W., A. Sittkus, H. Stockburger, and H. Sartorious, Large- scale atmospheric mixing

derived from meridional profiles of krypton 85, J. Geophys. Res. 88, 8574-8578, 1983.

-- --

- 75 -

Wofsy, S. C., Interactions of CH4 and CO in the earth’s atmosphere, Ann. Rev. Earth Planet.

Sci., 4, 441-469, 1976.

Wofsy, S. C., Temporal and latitudinal variations of stratospheric trace gases: a critical com-

parison between theory and experiment, J. Geophys. Res. 83, 364-378, 1978.

Yang, H. and K. K. Tung, Cross isentropic stratospheric-tropospheric exchange of mass and

water vapor, J. Geophys. Res., 101, 9413-9424, 1996.

Zimmerman, P. R., J. P. Greenburg, and C. E. Westberg, Measurements of atmospheric hydro-

carbons and biogenic emission fluxes in the Amazon boundary layer, J. Geophys. Res., 93,

1407-1416, 1988.

Zimmermann J, and D. Poppe, Nonlinear chemical couplings in the tropospheric NOx-HOx

gas-phase chemistry, J. Atmos. Chem., 17, 141-155, 1993.

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Figure captions.

Fig.1. Vertical profiles of NO. The heavy solid lines show the profile adopted for 3

regions: 30°N- 90°N; 30°N-90°S (based on data from 30°N-35°S), and the southern tropics

from 60°W to 45°E in August to October. Vertical profiles from smaller regions are taken from

Wang et al., [1998], where the regions are defined, with the addition of data from PEM-Tropics

for which the data were averaged over the following areas: Tropical S. Pacific (region 26, 5°N-

15°S, 170°E-130°W); Sub-tropical S. Pacific (region 27, 10°S-35°S, 170°E-145°W); Easter

Island (region 28, 10°S-35°S, 120°W-105°W); New Zealand (region 29, 35°-55°S, 170°E-

170°W); Antarctic (region 30, 55°-75°S, 170°E-170°W); S. Easter Island (region 31, 35°-55°S,

105°-115°W). The profiles from the sub-regions are averages over all NO points obtained with

solar zenith angle of less than 70°.

Fig. 2. Comparison of model results (lines) used to specify the global distribution of CO

with observations (symbols) at selected sites at the surface (a) and for the column (b).

Fig. 3. Concentration of OH (105mol cm−3) at the surface at 15°S vs. concentration of

isoprene in January, with NOx at 67 pptv (solid line) and in July, with NOx at 219 pptv (dashed

line).

Fig. 4. Vertical profiles of OH (105mol cm−3) in the tropics for the cloud with reflectivity

0.4 extending to 700 mb (dashed line) and to 300 mb (solid line). The optical depth of the cloud

is distributed uniformly with altitude from the cloud top to 900 mb.

Fig. 5. Cloud reflectivity for ISCCP clouds vs. cloud-top pressure for 1990 in the tropics

(left panel), in summer in the extratropics in the northern and southern hemisphere (middle and

right panels, respectively).

Fig. 6. Zonally and monthly averaged concentrations of OH in 105mol cm−3 for January,

April, July and October.

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Fig. 7. Distribution of OH at 700 mb in 105mol cm−3 for January and July (upper and

lower panels, respectively).

Fig. 8. Annually and globally averaged concentrations of OH in 105mol cm−3 for the

present distribution (solid line), for the OH from S90 (dashed line) and for OH computed with no

isoprene (dotted line). Concentrations were integrated with respect to the mass of air.

Fig. 9. Major chemical reactions affecting odd-H and HOx in the lower and middle tropo-

sphere.

Fig. 10. Fractions (%) of the total production of OH in reactions O(1D)+H2O→2OH,

NO+HO2->OH+NO2 , O3+HO2->OH+2O2 and their sum (left panel). Fractions of the total loss

in reactions CO+OH->H+CO2 , CH4+OH->CH3O2+H2O, combined H2O2+OH->HO2+H2O,

CH3OOH+OH->CH3O2+H2O and CH2O+OH->CO+HO2+H2O, and their sum (right panel).

Fig. 11. Relative change in OH (%) in response to changes by -50%,-25% (dotted lines),

+25%,+50% (dashed lines) at 15°S over ocean in January in O3 (a), H2O (b), NOx (c), and CO

(d); and over land in July in response to changes in O3 (e), H2O (f), NOx (g), and CO (i). Note

that results for July reflect the influence of biomass burning.

Fig. 12. Errors in predicting lifetimes of gases with respect to destruction by tropospheric

OH, with the rate constant expressed as Aexp(-B/T), by scaling the rate coefficient to that for

CH3CCl3 [Prather and Spivakovsky, 1990] at 277K (short-dashed line), at 272K (bold solid

line), at 270K (thin solid line) and at 267K (long-dashed line). These errors were computed

assuming uniform mixing ratio throughout the globe. Larger errors may be incurred for tracers

with distributions displaying steep gradients (see text).

Fig. 13. Evaluation of the computed distribution of OH using observations of various

tracers. Crosses signify a ratio of the computed mean to that implied by observations of the

given tracer (e.g., a ratio 1.1 means that the given constraint points to computed OH in the

specified region being too high by 10%). Solid lines denote the range of uncertainties of that

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ratio assuming that the mean OH implied by observations of tracers is known exactly. Uncer-

tainties associated with determining the mean OH implied by observations, are given with

respect to ratio=1 and are denoted by dashed lines. They include uncertainties in the magnitude

of sources (SR), absolute calibration (AC), strength of other sinks (OS and SS for ocean and stra-

tospheric sinks, respectively), and rate constant for reaction with OH (RC).

Fig. 14. Observed long-term trend in CH3CCl3 [Prinn et al., 1995] (dots) simulated using

(1) standard OH without the ocean sink (solid lines), and with tropospheric loss frequencies

reduced and increased by 25% (dashed lines).

Fig. 15. Observed long-term trend in HCFC-22 [Montzka et al., 1996] (circles), [Miller et

al., 1998] (triangles) simulated using (1) standard OH (solid lines) and (2) and (3): the standard

OH reduced and increased by 25% (dotted lines).

Fig. 16. Observed (circles) and simulated concentrations of CH2Cl2 using standard OH

and standard rate of interhemispheric mixing, D=180km (bold solid lines); standard OH and

increased interhemispheric mixing, D=250 km (thin solid lines); OH increased in the northern

hemisphere by 35% and decreased by 60% in the southern tropics (which led to a decrease of the

hemispheric mean of OH in southern hemisphere by 45%), and D=180 km (chain-dashed lines);

standard OH, D=180 km, with a source of 40 Gg of CH2Cl2 distributed uniformly over the

ocean surface (dotted lines). Comparisons are presented as 12-month running means (omitted if

monthly values are available for fewer than 10 month encompassing 6 previous and 5 following

months).

Fig. 17. Observed (circles) and simulated concentrations of C2Cl4 using standard rate of

interhemispheric mixing, D=180km, and standard OH (bold solid lines); standard OH, but the

rate constant for reaction with OH increased and reduced by 30% (short-dashed and dotted lines,

respectively); standard OH, recommended rate constant [DeMore et al., 1997], and a day-time

mean concentration of Cl of 1.25×104 mol cm−3 in the lowest 500 m over the oceans (long

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dashed lines). Comparisons are presented as 12-month running means (omitted if monthly

values are available for fewer than 10 month encompassing 6 previous and 5 following months).

Fig. 18. Latitudinal distribution of 85Kr from the Atlantic cruise in March 1983. The solid

line shows results for the standard model (D=180km), the dashed line shows results for the simu-

lation with increased interhemispheric mixing (D=250km).

Fig. 19. Latitudinal gradient of CFC-11 as represented by differences between concentra-

tions at ALE/GAGE stations and those at Tasmania (41°S, 145°E), averaged over 1980-1983.

Observations [Cunnold et al., 1994] are denoted by closed symbols: circles for CFC-11 (silicone

column), and diamonds for CFC-11 (Porasil column). The simulation is shown by open squares.

For both model and observations, monthly values for each site were obtained using a second-

order polynomial providing the least square fit to monthly means for the running-30-day medi-

ans, to filter out the influence of local pollution and to minimize the impact of missing data. The

distribution and history of emissions were taken from McCulloch et al. [1994] and Fisher [1994].

Fig. 20. Latitudinal gradient of CH3CCl3 as represented by differences between concentra-

tions at ALE/GAGE stations and those at Tasmania (41°S, 145°E), averaged over 1980-1983.

Observations [Prinn et al., 1995] are denoted by closed circles. Simulations are shown by open

symbols: squares for standard model and triangles for an interhemispheric ratio in OH of 50%,

north to south (pointed down) and south to north (pointed up). For both model and observations,

monthly values for each site were obtained using a second-order polynomial providing the least

square fit to monthly means for the running-30-day medians, to filter out the influence of local

pollution and to minimize the impact of missing data. The distribution and history of emissions

were taken from Midgley and McCulloch [1995].

Fig. 21. Annual cycle CH3CCl3 at Tasmania averaged over 1980-1991 (%). Relative resi-

duals from the long-term trend were computed in a manner described in S90. Triangles denote

observations [Prinn et al., 1995], with error bars corresponding one standard deviation from the

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mean. Dashed lines show simulations with concentrations of OH reduced and increased by 50%

south of 28°S; the thin solid line shows the simulation with aseasonal OH: concentrations of OH

were averaged over the year in each grid box.

Fig. 22. Observed and simulated relative seasonal variations of CH2Cl2 at Cape Grim (%).

Symbols show averages for each months of monthly means, relative to the yearly average, over

1994-1997. Error bars depict the full range of monthly values over this period. The simulation

with standard OH corresponds to the bold solid line, and those with concentrations of OH

reduced and increased by 50% south of 28°S are shown as dashed lines.

Fig. 23. Observed and simulated relative seasonal variations of C2Cl4 at Cape Grim (%).

Symbols show averages for each months of monthly means, relative to the yearly average, over

1994-1997. Error bars depict the full range of monthly values over this period. The simulation

with standard OH corresponds to the bold solid line, and those with concentrations of OH

reduced and increased by 50% south of 28°S are shown as dashed lines.

Fig. 24. Comparison of observed (bold solid line) and simulated relative seasonal varia-

tions of C2H6 at Harvard Forest [Goldstein et al., 1995]. The simulation with standard OH is

shown as a thin solid line, and those with standard OH reduced and increased by 50% north of

28°N as dotted lines. Observations and simulations are presented as a 30-day running 10%-

quantile relative to their yearly mean.

Fig. 25. Observed and simulated relative seasonal variations of CH2Cl2 in northern extra-

tropics (%). Symbols show averages for each month of monthly means, relative to the yearly

mean , over 1994-1996. Error bars depict the full range of monthly values over this period. The

simulation with standard OH is shown as a bold solid line, and that with concentrations of OH

increased by 50% north of 28°N is shown as dashed lines.

Fig. 26. Observed and simulated relative seasonal variations of C2Cl4 in northern extratro-

pics (%). Symbols show averages for each month of monthly means, relative to the yearly

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mean, over 1994-1996. Error bars depict the full range of monthly means over this period. The

simulation with standard OH is shown as a bold solid line, and that with concentrations of OH

increased by 50% north of 28°N is shown as dashed lines.

Fig. 27. Annual cycle of 14CO at Baring Head, New Zealand averaged over 1989-1991:

residuals with respect to the annual mean in % (upper panel), concentrations on absolute scale

(lower panel). Solid lines correspond to the simulation with the standard dynamics. Dotted lines

show the simulation with the rate of strat-trop exchange decreased by a factor of 2. Dashed lines

correspond to the simulation which in addition to reduced air-flux from the stratosphere,

included relocation of emissions that the model erroneously attributes to the troposphere (see

text). Long-dashed lines show the simulation including two previous corrections, and an addi-

tional diffusion of 6×1010cm2sec−1 south of 28°S in austral winter combined with vertical diffu-

sion of 8×105cm2sec−1 south of 60°S. The thin solid line in the lower panel shows results for

the simulation with no cosmic source. The standard distribution of OH is used in all simulations.

Observations (symbols) are from Brenninkmeijer et al., 1993. Error bars represent the full range

of monthly mean values over 1989-1991.

Fig. 28. Global distribution of excess 14CO2 observed in July 1966 (solid lines with closed

triangles). Simulations were initialized in October 1963 using observations taken from Johnston

[1989] along with the lower-boundary conditions. Solid lines with closed circles show the simu-

lation with the standard model. Dotted lines with open circles correspond to simulations with

the air flux through the 150 mb surface reduced by 25% and 75%. Solid lines with open circles

show the simulation with the air-flux reduced by 50%.

Fig. 29. Latitudinal distribution of 14CO at about 6-7 km altitude (upper panels) and at the

surface (lower panels) in January-February (left panels) and August (right panels). Observations

for 1990-1991 (diamonds) are from Mak et al. [1992], and for 1992 (squares) and 1993 (trian-

gles) are from Mak et al. [1994]. Note that the cosmic source was approximately constant in

1990-1991, increasing by 8% in 1992 and additional 14% in 1993. See Figure 27 for the

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description of simulations and designation of line types. To facilitate a comparison with obser-

vations, model results south of 32°N are shown for 1990-1991, and north of 32°N for 1993. We

also included observations in 1977-1978 from Volz et al., [1981]; the cosmic source for that

period was comparable to that in 1992-1993.

Fig. 30. Vertical gradients of 14CO in 1990-1991 at southern midlatitudes in January-

February (left panel) and August (right panel). See Figure 27 for the description of simulations

and designation of line types. Observations (symbols) at the surface are from Brenninkmeijer et

al. [1993] and aloft from Mak et al. [1992].

Fig. 31. Seasonal variations of 14CO in 1977-1978 at northern midlatitudes: residuals with

respect to the annual mean (a), on absolute scale (b). See caption Figure 27 for the description

of simulations and designation of line types. Observations are from Volz et al. [1981].

Fig. 32. Annually averaged latitudinal distribution of CH3CCl3 in excess of concentrations

at Tasmania relative to the surface mean concentration in 1999 (upper panels) and 2002 (lower

panels) with zero emissions beginning from January 1997 (left panels) and 30 Gg y−1 (right

panels). Zonal means for simulations with the standard OH are shown as solid lines, and those

with an interhemispheric ratio in OH of 1.5, north to south, and south to north, by dotted and

dashed lines, respectively. The lifetime of CH3CCl3 in all simulations is 5.0 years. The ocean

sink is not included. Results for selected NOAA and ALE/GAGE stations (at South Pole,

Tasmania, Samoa, Barbados, Alaska, Ireland) are shown as open circles.

Fig. 33. Latitudinal distribution of CH3CCl3 in excess of concentrations at Tasmania rela-

tive to the surface mean concentration in February 1999 (upper panels) and August 1999 (lower

panels) with zero emissions beginning from January 1997 (left panels) and 30 Gg y−1 (right

panels). Zonal means for simulations with the standard OH are shown as solid lines, and those

with an interhemispheric ratio in OH of 1.5, north to south, and south to north, by dotted and

dashed lines, respectively. The lifetime of CH3CCl3 in all simulations is 5.0 years. The ocean

sink was not included. Results for selected NOAA and ALE/GAGE stations (at South Pole,

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Tasmania, Samoa, Barbados, Alaska, Ireland) are shown as open circles.

Fig. 34. Zonally and annually averaged lifetime of CH3CCl3 (y) computed for the standard

OH.

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