NASA Reference Publication 1010 

-(NASA-RP-1lOO) CHLOROFLU6ROIETHANES AND THE
 

STRATOSPHERE (NASA) H
BC A12/MF A01
 

CSCL 04A
 

Chlorofluoromethanes
 

and the Stratosphere
 

N77-31694i
 

THRU 

N77-31699
 

Unclas 

H1/46 47596
 

August 1977 

REPRODUCEI) BY 

NATIONAL TECHNICAL 

INFORMATION SERVICE 

U.S. DEPARTMEN'T OF COMMERCE 

SPRINGFIELD, VA. 22161 

NASA
 
1. Report N! 2. Goverment Accosson Njo, 3. Recipient's Cotalog No. iP-1010
 

4. Title and LStitle 5. Report pate 

ChlIorofluoromethanes in the Stratosphere September 1, 1977 

6- r~mn .gOrganization Code 

624 

7. Author(s) 8. Performing Orgonization Report HO. 

R. D. Hudson, et al. G-7725 

9. Performig OrganizationName and Addess 10. Work Unit No 

i98-20-01 Goddard Space Flight Center 

11. Contract or Groat No. Greenbelt, Maryland 20771 

13. Type of Repat a.d Period Covered 

12. Sponsormng Agency Name ard Address Reference Publication 

National Aeronautics and Space Administration 

Washington, D.C. 20546 14. Sponsoring Ailncy Code
 

A-12
 

15. Supplementary Notes 

16. Abstract 

This document reports the conclusions of a workshop held by the National 

Aeronautics and Space Administration to assess the current knowledge of 

the impact of chlorofluoromethane release in the troposphere on stratospheric ozone concentrations. The document is divided into five sections: 

Laboratory Measurements, Ozone Measurements and Trends, Minor Species 

and Aerosol Measurements, One-Dimensional Modeling, and Multidimensiona Modeling. 

17. Key Words (Selected by Author(s)}) 1 Distribution Statement 

Halocarbons, Stratosphere, Ozone, Unclassified - Unlimited 

Chlorofluoromethanes STAR 23 Chemistry and Materials 

1M Security Classif. (of this repor) 20. Security Classif. (of this V 

Unclassified Unclassified 

-For sale by the National Technical Information Service, Springfield, 

I 
NASA Reference Publication 1010
 

Chlorofluoromethanes
 

and the Stratosphere
 

/ 

Robert D. Hudson, Edlitor 

Goddard Space Flight Center 

Greenbelt, Maryland 

National Aeronautics 

and Space Administration 

Scientific and Technical 

Information Office 

1977 

iW 
For sale by the National Technical Information Service 

Springfield, Virginia 22151 
PREFACE
 

In 1976, the National Aeronautics and Space Administration (NASA) 

made a commitment to provide, by September 1977, an assessment of 

the effect on the stratospheric ozone of the release of chlorofluoromethanes into the troposphere. It was recognized at that time that the 

NASA report would closely follow a report prepared by the National 

Academy of Sciences, "Halocarbons: Effects on Stratospheric Ozone," 

dated September 1976. Therefore, the NAS report was used as the 

starting point in the preparation of the NASA report. 

To gather the latest information for such a technical update, the Goddard 

Space Flight Center organized a workshop at the Airlie House, Warrenton, 

Virginia, during the week of January 10, 1977. The workshop was 

centered around five working groups: a Laboratory Measurements working group chaired by Dr. W. B. DeMore of the Jet Propulsion Laboratory; 

an Ozone Measurements and Trends working group chaired by Prof. I. 

London of the University of Colorado; a Minor Species and Aerosol 

Measurements working group chaired by Prof. R. Megill of Utah State 

University; a One-Dimensional Modeling working group chaired by 

Dr. R. S. Stolarski of the Goddard Space Flight Center; and a Multidimensional Modeling working group chaired by Prof. M. Geler of the 

Ur Tversity of Illinois. The overall chairman for the workshop and for the 

production of the workshop report was Dr. R. D. Hudson of the Goddard 

Space Flight Center. 

Each working group was further subdivided, and prominent scientists 

who are active and expert in stratospheric studies were asked to prepare 

position papers to cover selected topics of research. In addition to preparing position papers, in October 1976, modeling groups were approached 

to "run" their models using a standard set of input data supplied by the 

Goddard Space Flight Center. Specific output data, as well as the predicted ozone changes, were requested so that realistic model intercomparisons could be made. 

In all, over sixty scientists attended the workshop, representing most of 

the institutions engaged in stratospheric research in the United States, 

Canada, and Europe. A complete list of the position-paper authors and 

participants, with working group assignments, is given in the Appendix. 

iii 
Each working group prepared a summary document, and these have been 

assembled into this report. The report has been updated since January 

1977 by soliciting written comments from the participants and as a result 

of further meetings of the Laboratory Measurements and the OneDimensional Modeling working groups in June and July 1977. Thus, the 

report should represent current knowledge as of June 1977. 

Goddard Space Flight Center and the National Aeronautics and Space 

Administration express their thanks to the Working Group Chairmen, the 

position-paper authors, and the participants for the time and efforts that 

they contributed to the preparation ofthis report. 

iv 
CONTENTS 

Page 

PREFACE .. F.A.C.E.. . 

CHAPTER 1 - LABORATORY MEASUREMENTS .... 1
 

INTRODUCTION . . 1
 

Nitrogen Oxides .
 

Hydrogen.Oxides ....... .............. 2
 

Chlorine Oxides ............... 3
 

Bromine Reactions 

Photochemical Processes ....... 55..... 


...... ............. 4
 

Fluorine Reactions 5.............5
 

DEFINITIONS ....... ................ 6
 

Units ........................... 6
 

Reliability Estimates ..... ............ 6
 

Rate Parameters 7..............7
 

SUMMARY OF RATE-CONSTANT RECOMMENDATIONS 7
 

Notes to Table I ..... .............. .14
 

CROSS SECTIONS AND QUANTUM YIELDS ...... 25
 

Introduction ........... ......... 25
 

Discussion of Photodissociation Reactions ..... 26
 

HETEROGENEOUS CHEMISTRY . .......... 45
 

Removal of Chlorofluoromethanes (CFM's) ..... ... 46
 

Removal of Cl ..................... 47
 

Catalytic Processes..... ............. .. 48
 

CHAPTER 2 - MEASUREMENTS AND ANALYSIS OF
 

OZONE VARIATIONS .......... 51
 

INTRODUCTION ............ .... 51
 

MEASUREMENTS . ............... 52
 

Total Ozone ................ 52
 

Ozone Vertical Distribution .......... 57
 

v 
CONTENTS (Continued) 

Page 

ANALYSIS OF OZONE VARIATIONS ........ 
 67 

Data Availability ..... .............. .. 67
 

Long-Tern Variations ............ 71
 

Some Inferences on Photochemical Models Derived
 

from Ozone-Variation Observations ......
 80 

COMMENTS ON THE ANALYSIS OF OZONE DATA 82 

Distinguishing Between Natural and Anthropogenic
 

Trends ... ............. 82
 

Statistical Analysis Applied to Present Data Sets 84
 

Threshold of Detectability ....... 84
 

SUMMARY AND CONCLUSIONS ............ 86
 

CHAPTER 3 - MINOR SPECIES AND AEROSOLS ...... 89
 

INTRODUCTION ............... 89
 

SUMMARY OF ATMOSPHERIC MEASUREMENTS . 92
 

SOURCE SPECIES ... ............. 92
 

SINK TERMS .............. 95
 

RADICAL SPECIES .............. 103
 

SOLAR FLUX AND O/O 3 RATIO ......... 115
 

TRANSPORT MEASUREMENTS ........ 121
 

Winds and Dynamics .......... 122
 

Aerosol Measurements ..... ............ .125
 

RECOMMENDATIONS ..... ............. .128
 

Measurements. ............... 128
 

Hydrogen Species ............. 129
 

Nitrogen Species .............. 130
 

Halogen Species 1..............31
 

CHAPTER 4 - ONE-DIMENSIONAL MODELS .......
 133
 

INTRODUCTION ...............
 133 

PARAMETRIC SIMULATION OF NET VERTICAL
 

TRANSPORT IN 1-D MODELS ............ 134
 

vi 
CONTENTS (Continued) 

Page 

Conceptual Basis for the Introduction of Kz . .... 135
 

Chemical Tracers ........ .......... .. 136
 

Radionuclide Tracers .... .......... 138
 

GCM Tracer Experiments .... ........... .. 139
 

DIURNAL AVERAGING EFFECTS ... ....... ... 139
 

MULTIPLE-SCATTERING EFFECTS .... ...... .144
 

Photodissociation Coefficients .. ....... . . . 144
 

Species Concentration Profiles ... ........ .148
 

Model Sensitivity ... ............... .152
 

Summary ...... ................ .152
 

COMPARISON OF THEORY AND
 

MEASUREMENT: HO .... ............ 152
 x
 

COMPARISON OF THEORY AND
 

MEASUREMENT: NOx ..... ............ 157
 

COMPARISON OF THEORY AND
 

MEASUREMENT: CIOX .... ............ .168
 

Introduction ...... ............. 168
 

Sources of Stratospheric Odd Chlorine ...... 169
 

Measurements and Theoretical Predictions
 

for CIX Species ..... ............. .. 172
 

TROPOSPHERIC AND STRATOSPHERIC SINKS
 

FOR CHLOROFLUOROMETHANES .......... .178
 

Introduction ....... ............. .. 178
 

Removal by Oceans ... . ............ 180
 

Photodissociation in the Troposphere ........ .181
 

Reaction with OH Radicals in Troposphere ..... ... 181
 

Removal by Pyrolysis . ..... ..... .182
 

Ion-Molecule Reactions .... ........... .183
 

Active Removal in Stratosphere ...... . . . . 184
 

Conclusion ....... ................ .184
 

EFFECT OF TEMPERATURE FEEDBACK
 

MECHANISM ....... ................ .. 185
 

UNCERTAINTIES IN THE PREDICTIONS OF 1-D
 

STRATOSPHERIC MODELS .. .......... 186
 

Introduction . . . ........... .. 186
 

Reaction Rates ... ........... 186
 

vii 
CONTENTS (Continued) 

Page 

CONCLUSIONS ....... 189
 

CHAPTER 5 - MULTIDIMENSIONAL CFM-O MODELING. . . 197
 

INTRODUCTION ..... .............. 197
 

RADIATIVE-DYNAMICAL]PHOTOCHEMICAL
 

FEEDBACK MECHANISMS .......... . 199
 

Background ......... ........ .199
 

Feedback Mechanisms ..... ............ .200
 

Concluding Remarks .... ............ 205
 

TWO-DIMENSIONAL MODELING .......... 205
 

Types of Models .............. 206
 

Chemical Model ..... .............. .207
 

Atmospheric Transport Prescriptions for
 

2-D Models ....... .............. .. 208
 

Chemistry and Transport ........... 210
 

Boundary Conditions .... ............ .. 210
 

Applications ....... ............... 211
 

Summary .... ............ -.... 213
 

THREE-DIMENSIONAL MODELING ... ........ ... 214
 

TRANSPORT PARAMETERIZATIONS IN 3-D
 

MODELING OF THE STRATOSPHERE ......... .. 218
 

Subgrid-Scale Mixing .... ............ .219
 

Vertical Transport Across the Tropopause ..... 220
 

OBSERVATIONAL NEEDS OF MULTIDIMENSIONAL.
 

MODELS OF THE STRATOSPHERE .......... .. 221
 

CONCLUSIONS ....... ............. .226
 

Assessment Uncertainty Due to Transport ...... .. 226
 

Contributions of Multidimensional Models ...... .. 227
 

Observational Needs .... ............. ... 228
 

REFERENCES ....... ........... .... 231
 

APPENDIX - CHLOROFLUOROMETHANE WORKSHOP
 

ATTENDEES ..... ............. .. 261
 

viii 
N77w31695
 

CHAPTER I 

LABORATORY MEASUREMENTS 

INTRODUCTION 

This section presents a compilation of chemical and photochemical data 

that is relevant to stratospheric modeling. There are three broad categories of data: (1) rate constants for chemical reactions,including (when 

possible) temperature and pressure dependencies along with product distributions; (2) absorption cross sections, photodissociation quantum yield, 

and photolysis products; and (3) heterogeneous chemical processes. 

The complexity of stratospheric chemistry, combined with the need for 

accurate model predictions, has resulted in stringent demands for reliable 

laboratory data in the categories noted above. In many instances, adequate 

measurements are still not available, partly because measurement programs 

are not complete and partly because satisfactory measurement techniques 

are either unavailable or have become available only very recently. Moreover, agreement among measurements by different techniques has occasionally been unsatisfactory, as the detailed discussion in subsequent 

subsections will show. 

The recommendations are based entirely on laboratory measurements, 

with occasional extrapolations using theory. An example of the latter is 

the pressure falloff in the rate constant for C1NO 3 formation at pressures 

relevant to the stratosphere. (This question is discussed in detail in subsequent subsections.) No "tuning" of rate data to fit stratospheric measurements was employed. All laboratory data judged to be reliable were 

considered, including measurements of absolute rate constants and ratios 

of rate constants. 

NITROGEN OXIDES 

The reactions forming and destroying NO, N0 2 , NO3 , and HNO3 have 

been part of modern atmospheric modeling since the beginning of the consideration of the environmental effects of supersonic transport aircraft. As 

a result, considerable attention has been paid to the measurement of the 

rate constants of these reactions. For many of them, the reliabilities are 

near the limit of the measurement art (perhaps 25 to 40 percent agreement 

betweenlaboratories). Also, for many of thbm, these reliabilities appear to 

be sufficient for modeling purposes as judged from sensitivity calculations. 

1
 
2 	~4 ~CHLOROFLUOROMETIIANE WORKSHOP REPORT 

However, there are reactions for which the data base is weak and for which 

difficult, and occasionally arbitrary, decisions have necessarily been made. 

These reactions, which become candidates for further study, are listed 

below; the problems are discussed in the notes that accompany table 1. 

* 	 OH + NO 2 + (N2 andO0 2) -+ HN0 3 + (N2 andO ) atlowtemperatures and high pressures-One set of such measurements exists. 

It is difficult to correlate with theory. Verification is needed. 

" 	 HO2 + NO2 (+M) -> HO2N0'2 (+M) under all stratospheric conditions-Peroxynitric acid is a molecule that has not yet been used in 

atmospheric models but must be considered. 

* 	 Photochemistry of HO2NO2 and its reactions with 0, OH, and other 

free radicals-The chemistry of HO2N2 must be known before this 

species can be introduced in the models. 

" N + NO 2 -) N2 0 + 0andN + 0 3 -- NO + 02 todetermine 

the temperature dependences that have been estimated in the absence 

of data. 

resolved 

(in 

tainty 

of 

The 

these 

two 

reaction-rate 

of 

laboratories) 

problems. 

sets. 

either. 

Experiments 

It 

data 

Measurements 

has 

differ 

for 

been 

0('D) 

systematically 

to 

decided 

resolve 

are 

using 

reasonably 

to 

this 

twd 

base 

by 

incompatibility 

different 

the 

more 

reliable 

recommendations 

than 

analytical 

but 

the 

would 

have 

known 

techniques 

some 

improve 

on 

uncerone 

unthe reliability of the recommendations. In addition, the reactions of 0 (D) 

with chlorocarbons and halogenated derivatives of formaldehyde require 

more study. The products of the chemical reactions have not been determined, and it is not known whether the observed rate constants contain an 

appreciable contribution from physical quenching. 

HYDROGENOXI DES 

There are only two HO, reactions that may be known well enough to require no additional study: 

* 	H+ 0 2 +M-+H0 2 +M 

SOH + CH4 -+ CH3 + H2 0 
LABORATORY MEASUREMENTS 3 

The characteristic situation for the remaining reactions (especially atom/ 

radical and radical/radical) are a reasonably well-known room temperature 

rate constant, but there is no temperature dependence data. These require further study, preferably by direct techniques. In view of the sensitivity study carried out in conjunction with this assessment, top priority 

should be focused on the following reactions: 

* OH + HO > IH 
 2 -- 20 +0 2

SH02 + HO2 -H202 + 02 

SHO2 + NO- NO2 + OH 

SHO2 +03 "OH + 20 2 

* OH + 03 -HO2 +0 2 

. H02 + 0 -OH + 02 

Existing data for the first three reactions are mutually inconsistent, and, 

thus, further work is required on all three. The rate constants recommended here are an adjusted set, as discussed in note BI to table 1.The 

remaining three reactions contribute significantly to the total uncertainty 

of the stratospheric models. 

CHLORINE OXIDES 

Although many ClOy reactions are listed in table 1, Summary of Recommended Rate Constants, they are not of equal importance. Accurate 

models require precise kinetic data for reactions 44 through 50 because the 

catalytic efficiency is strongly dependent on the absolute magnitude of 

these reaction-rate constants. Although it is essential to include reactions 

51 through 77 in any complete photochemical model calculation, these 

reactions are of less importance, and, thus, a lower precision can be tolerated in the kinetic-rate data for these processes. Whether reaction 49 

plays an important role in stratospheric chemistry is highly dependent on 

the [HO2] profile that is governed by HOx reactions. Therefore, whereas 

an uncertainty of 20 percent in knowledge of the values of k44 through ks0 

may be significant, an uncertainty by a factor of 2 or more in the values of 

k., through k77 may not be particularly important for calculating the magnitude of the ozone depletion. However, accurate rate-constant data may 

be required for a larger set of reactions if an understanding is required of 
4 CHLOROPLUOROMETHANE WORKSHOP REPORT 

the concentration of some minor chlorine containing species in the 

atmosphere (e.g., HOCI). 

One reaction that had not been previously studied at all (C10 + NO2 + M 

CIONO 2 + M) has now been the subject of extensive investigations at low 

pressures (equal to or less than 6 tort), but none at the relevant stratospheric pressures (approximately 50 torr). The recommended rate constant 

for ClONO2 formation is based on the evaluation of Zahniser et al. (1977). 

Data for that evaluation were obtained from three low-pressure, dischargeflow, temperature-dependent studies (Leu et al., 1977; Zahniser and 

Kaufman, 1977*; and Birks, 1977). In the absence of experimental data 

for pressure falloff behavior, Zahniser et al. (1977) fitted the pressure 

dependence to a theoretical Rfce-Ramsperger-Kassel-Marcus (RRKM) 

model, as described by Smith and Golden (1977). 

To date, there have been no quantitative determinations of the products of 

the 'C10 + NO2 + M reaction, and, at present,it is not possible to rule out 

the existence of two primary reaction channels. Consequently, because this 

reaction is of importance, it must be studied experimentally within the 

relevant pressure regime with a quantitative determination of the reaction 

products. Additional studies of the NO + C10 reaction and, to a lesser 

extent, the 0 + C10 reaction are required. 

Reactions 78 through 88 have been included in this compilation for the 

sake of completeness so that tropospheric lifetimes can be calculated for 

these compounds. 

BROMINE REACTIONS 

Compared to current knowledge of ClO reaction-rate data, there is a 

dearth of precise kinetic information for BrOx reactions. However, there is 

probably sufficient information to perform a meaningful sensitivity calculation. The basic set of key reactions in the BrOx system is similar to that 

in the C1O x system, with the following exceptions: (1) the Br + H2 and 

Br + CH4 reactions-are highly endothermic and are too slow to be significant, and (2) the radicayradical reactions (reactions 91 and 92) play an 

important role. The reactions that require the most immediate attention, 

ifthe BrOx system is'to be modeled with any real accuracy, are 95, 90, and 

93. Although there is no evidence for reaction 93, if it had a significant 

*M- S. Zahniser and F. Kaufman, University-of Pittsburgh, manuscript in preparation, 

1977. 
LABORATORY MEASUREMENTS 5 

rate constant (> 10" ? cm 3 sl), it would play a major role in the destruction of odd oxygen. The BrO x system can be generalized by stating that 

the reactivity of atomic bromine is less than that of atomic chlorine, 

whereas the reactivity of BrO is similar to, or greater than, that of CIO. 

All BrOx reactions require additional study. 

FLUORINE REACTIONS 

Although both atomic fluorine and the FO radical are extremely reactive 

species, the catalytic efficiency of FOx for the destruction of odd oxygen 

is very low due to the unreactive nature of the HF molecule. Atomic 

fluorine is highly reactive toward any hydrogen containing molecule and 

abstracts a hydrogen atom to form HF. HF does not react with the OH 

radical or the 0 (SP) atom and does not photolyse in the stratosphere. 

Therefore, as FO. quickly forms HF, the catalytic chain length is very 

short. However, this simple approach neglects the role of F0 2 radicals 

that are produced in the F + 02 + M reaction. 

PHOTOCHEMICAL PROCESSES 

Among the photolytic reactions mentioned later in table 3, the following 

are singled out as requiring further study: 

* 	 03 + hw - O('D) + 0.-The quantum yield values have been measured relative to the value below 300 nanometers, which is found 

to be nearly constant and assumed to be unity. There are, however, 

indications that the quantum yield is less than unity below 300 

nanometers. 

* 	 CIONO2 + hv - Pro ducts-The temperature dependence of the cross 

sections above 300 nanometers should be measured, and the identity 

of the photolytic fragments needs to be established. 

* 	 N2 05 + hp -+ Products-The quantum yield for photodissociation 

and the identity of the products have not been determined yet. 

* 	 The mechanism of photo-oxidation of the chlorofluoromethanes is 

not clearly understood, and the chemistry of the intermediate peroxy 

radicals should be studied further. In addition, the quantum yields 

for the photolysis of CF0 and CFC10 should be measured. 
6 CHLOROFLUOROMETHANE WORKSHOP REPORT 

DEFINITIONS 

UNITS 

Definitions are provided in the subsections that follow for use with table 1. 

The rate constants are given in terms consistent with concentrations in 

molecules per cubic centimeter and time in seconds. Thus, for first-, 1 3 t second-, and third-order reactions, the units of k are s ,cm molecule "

"1 2 " S ,and cm 6 molecule " s ,respectively. 

RELIABILITY ESTIMATES 

Reliability estimates are listed in table I under the column headed "±log k 

(230 K)" and apply at 230 K. They allow for any extrapolation of data 

from higher temperatures. The reliability stated in table 1 is a factor to be 

used to find the limits within which k probably lies. It is used in either of 

the following equivalent forms: 

log10 kL = loglok 0 - log 0

,

a lOgl ok u = 1Ogl ok o + loglo 

kL= k/a ku k
 

where k.is the central value reported for the rate constant, andloga is the 

entry in the ±logk (230 K) column. 

These rellabilities are estimates of overall accuracy. They include consideration of precision of measurement, performance of the methods on other 

reactions, agreement of measurements made in different laboratories, interrelationships among values of k for several reactions, and reasonableness in 

terms of theory. The recorded reliability, log l a,has been scaled to be 

roughly analogous to one standard deviation, a, because it is used as a 

measure of precision. This means that, for these reliabilities, the odds are 

5 to 3 that k will lie within the calculated upper and lower limits. 

Twice the value of logloor may be used to calculate limits analogous to the 

95-percent confidence interval (i.e., odds of 19 to 1 that k will lie between 

the limits). These definitions are to be taken as descriptions of ranges and 

not as implying a normal distribution of errors of measurement. There are 

usually two few sets of measurements to permit the distribution to be 

studied. When there are many measurements from several laboratories, 

there is no clear evidence that the distribution is normal. 
LABORATORY MEASUREMENTS 7 

RATE PARAMETERS 

The rate constants given in table 1 are expressed, almost without exception, 

in the form k = A exp (-BIT). This is the simple Arrhenius form, with A 

being the "pre-exponential factor" and B the "activation temperature." 

Here, B = E*/R, where E* is the apparent activation energy, and R is the 

gas constant. 

The values of A and B reported here have been developed to reproduce the 

experimental rate data that have been accepted as reliable. Both A and B 

should be considered to be empirical fitting parameters, suitable for deriving a value of k for the limited temperature range 200 to 300 K. 

On the basis of theoretical models and experience, kineticists expect values 

of A for particular classes of reactions to lie within fairly well-defined 

ranges. These ranges correlate with the expected structures of the transition states. Many of the values of A reported here meet these general 

criteria. However, some do not, and these are usually lower than expected. 

This is a warning that either the transition state or mechanism is not understood or that there are systematic errors in the trend of k with temperature. 

When faced with an unexplainable discrepancy between observed and 

expected parameters, an empirical fit to the data has usually been favored 

and a more generous uncertainty assigned than would be indicated by the 

experiments alone. 

SUMMARY OF RATE-CONSTANT RECOMMENDATIONS 

A summary list of recommended rate constants is given in table 1. Recommendations with explanatory notes designated "A," followed by a numeral, 

are derived from the position paper prepared for this assessment by D. 

Garvin and R.Hampson of the National Bureau of Standards (NBS). Similarly, reactions with notes designated by "B" and "C" are from position 

papers prepared by J. Margitan of the University of Michigan and R. Watson 

of the Jet Propulsion Laboratory, respectively. Reactions with notes designated by "D" represent recommendations made by the working group. 

Table 2 is an analytical expression for the rate constant of the reaction 

OH + NO2 + M - HONO2 + M,which is not suitable for representation in 

the Arrhenius form. 
Reaction 

1. 03 +NO - NO2 +02 

2. O+NO2 -* NO+O 2 

3. N+0 2 - NO+O 

4. 

5. 

6. 

7. 

8, 

9. 

1o. 


1I. 


12. 

13. 

14. 

15, 

16. 

17, 


18. 

19. 

20, 

'Notes 

N+NO - N2 +0 

OH+-NO 2 (0M) ' HNO 3 

OH+HNO3 -* H0 +NO 

N+NO2 - N20 +O 

NO+O+M - NO2 +M 

NO2 +HO 2 + N2 -+ HO2 No 2 

N+O 3 - NO+O 2 

NO 2 +03 "+NO3 +0 2 

Q(1D)+N2 0- N +0 2 

O(ID)+N0 -* NO+-NO 

O('D)+H 2 0 --QH+OH 

0(ID)+CH4 - OH+CH3 

O('D)+CH4 "* H2 +CHO 

O(D)+N 2 +M - N2 O+M 

O('D)+H 2 - OH+H 

0( 1 D)+N 2 - O+N 2 

0('D)+0 2 0+0 2 

appear in the text. 

Table I
 

Summary of Recommended Rate Constants
 

(+M) 


+N 2 

k 

2.1 X 10"2 exp (-1450/) 

"12
 9.1X 10 

" 5.5 X 10 2 exp (3220/T) 

"
 8.2X 101 exp(.410T) 

See table 2 

gX 10-14 

"
 2X 10'1 exp(-800/T) 


1.55X 10-32 exp(+584/T) 


See note 


2 X 10'1 exp (-1070/T) 


1.2X 10'13 exp(-2450/T) 
 "
 5.5X 1011 

5.5X 10"' 

23X 1Q0

" 

1.3X 1010 

"
 lAX t0­"
 3.5X 103 

9,9X 10-9 

2.0X 10'11 exp(+107/T) -
 2.9X 10 1 exp (+67/T) 

±log k (230 K) 

0.08 

0.03 

0.1 

0.1 

0.1 

005 

0.15 

0.05 

-

0.15 

0.04 

0.1 

0.1 

0.1 


0.05 

0.1 

0.5 

0.05 


0.05 


0.! 

Note* 

Al 

A2 O 

A3 o 

A4 

A5 o 

A6 a 

A7 

AS 

All 

A9 

A10 

A12 

A12 

A12 

A12 

A12 

A12 

A1I2 

A12 

A12 
Table 1 (Continued) 

Reaction k log 

k(230 K) Note* 

21. O(D)+ 0 3 -+ 02 +02 1.2X 10.10 0.1 A12 

22. 0(D)+03 -3 02+0+0 1.2X 10-"0 0.1 A12 

23. 0(D)+HC 

- OH+C1 IAX 10-

'

0 0,05 A12 

24. 0 ('D)+ CFCI3 

- products 2.3 X ! 10 0 0.1 A12 :b. 

25. 0(1D)+CF7.CI 2 

- products 2.0X 1010 01 A12 C 

26. 0( 1 D)+CCI2 0 

- products 1.7X 1IO 0.1 A12 

27. 0 ('D) + CFCIO 

- products 1.0 X 10"10 0.1 A12 o 

28, 0 ('D) + CF20 products 4.5 X 10"' 0.1 A12 

29. OH+HO2 H-2 0+02 3X 101 0.25 BI 

30, HO 2 +HO 02+O + 2.5X 10-12 0.3 BI 

31. NO+HO 2 NO2 +OH 2 8X 10 "

1 2-1..45 2 1+0.15 DI C 

32. HO 2 +0 3 OH+202 723X 101 4 exp(-1275/T) 0.3 B2 

33. OH+O 

z

3- HO 2 +O0 1 1.5 X 10"1 exp (-1000/T) 0.15 133 

34. O+OH -­

02 +H 4.2X 10 "1 ' 015 134 

35. O+HO -4OH+O 2 3.5X 10 " lt 0.15 B4 

36. 0+H

20"2 OH+H0 2 275X 10-12 exp(-2125/T) 0.15 15 

37. H+O 2 +M - HO 2 +M 2.1 X 10 " 3 2 cxp(+290/T) 0.1 136 

38. H+0 3 

- OH+O 2 1.2X 10 " 10 exp(-560/T) 0.15 137 

39. OH+OH -* H 2 0+O I X 10 " ' exp(-550/T) 0.2 u8 

*Notes appear in the text. 

'C 
Table 1 (Continued) 

Reaction k ±IoS k (230 K) Note* 

40. OH+OH+M -+ H202 +M 1.25X 1032 exp(+900/T) 0.J5 B9 

41 OH+1202 M2 0 + HO2 1X 10 " l exp (-750/T) 0,15 BIO 

42, OH+CO CO 2 +H 1AX 10" 3 0A BI1 0 

43. OH+CE4 H 2 0+CH3 2.35X 0 

12 exp (-1710/T) 011 B6 0 

44. CI+0 3 " CIO +02 2.7X 10" 1 exp(-2571T) 0.06 C1,2 

45. 0+ CIO 

- CI + 0 2 7.7 X 10 " exp (-1301T) 011 01,3 0 

46, NO+CIO - CI+N0 2 10X 10'-exp(+200/T) 0,15 C4 

47. 

48. 

OH +HMCI 

CI + CH 4 

- CI+H 0 2 

-+ [CI+CH3 

3.0X 10 " ' 

72 X 10 12 

exp(-425/T) 

exp (-1260/T) 

0.05 

'+0,06 

CI,5 r 

1-0.18 Ci,6 > 

49. CI+HO2 -HCI+O2 3X 10"11 03 CI,7 

50. CIO+NO +N 2 -* CIONO 2+N 3.3X 10-2T- ' 

- 4 

1 + 8.7 X 109 T 0 6 -'" 

1+015 

-0.06 

CS 

51. 0 + CIONO 2 - products 3.0 X 10

1

2 exp (-8087T) 0.15 CI,9 

52. OH + CION0 2 "+ products 1.2 X 1012 exp (-333/T) 0,18 Cl, 10 

53. Ci+ CiONO 2 -* products 1.68 X 10 2 exp (.607/T) 0.3 Cli 

54. 

55. 

0+MCI 

CI+0OH 

- CI+0OH - 0+HCI 

1.14X 10"' exp(-3370/T) 10X 10 "

11 exp(-2970T) 

013 

0,13 

C1,12 

C13 0 

56. CI + H2 " 1I + H 3.5 X 10 11 exp (-2290/T) 0,06 C1, 14 

57. CI+H20 2 -+ HCI+1H0 2 

- HOCI +Oi(?) 1.7X 10-2 exp (384/T) 0,18 C!, 15 

*Notes appear in the text. 
Table I (Continued) 

Reaction 	 k log k (230 K) Note* 

58. CI+H4N0 3 - HCI+N0 3 	 1.0X 10" I exp(-2170/T) 03 C1,16 "
 59. CI+NO2 +M -CIN02 +M 7.2 X 10-	 0.3 C1 

60. CI+NO4N 2 -* CINO+N 2 	 1.7X 10,

32 exp(+530/T) 03 C1 

61. CI+CINO -* NO+CI 2 30X 10-	 0.1 C1 

3 62. 	 01+O2 +M " C10+M 1.7X 1XT 3+1.0 

€-0.5 CI, 17 

63. CIOO+M -* CI+o 2 +M 5.8X I-9 exp(-35801T) 	 .+1.0

1-0.5 C1, 17 

"
 K1q = k62k63 2.95 X 1025 exp (+3580/T) 1.0 C1, 17 

64a. CI+COO -0. CIZ + 02 1.6X 10-1 +03 2 =-1.0 CI, 18 

64b. CI +CIO0 - CIO + CIO 1.1 X 1011 	 +0 5 

1-1.0 C1,]18 

65. CIOO + H02 . HCI + 202 No recommendation 	 C19 

66. CIO+HO H02-	 HeCl +02 	 2X 10.13 1 0 C20

" 67. CIO +CH4 "- products 	 <I X 1012 exp (-3700/T) 20 C1,21 

" 68. CIO+ H2 ->products 	 <1 ) 1012 exp (4800/T) 2.0 C1,21 " 69. CIO+CO - products 	 <1 X 1012 exp(-3700/T) 20 C1,21 "
 70. CIO+N 2 0 "* products <I1 X 1012 cxp(4260/T) 2.0 C1,21 

*Notes appear inthe text 
Table 1 (Continued) 

Reaction k ±Iog k (230 K) Note* 

71a. CIO+BrO 

- Br+OCIO 6.7X i~y 2 0.18 Cl 

71b. 

72a. 

CIO + BrO 

CIO+ CIO 

- Br + CIOO 

"+ CI + GOO 

6.7 X 10-12 

I X 10 " 2 exp (-I1238/T) 

0.18 

0.I 

CI 

C', 22O 0 

72b. CIO+CIO "+ CI2 +02 5X 1013 exp(.1238/T) 0.1 C1,22 

72c. CIO + CIO OCIO + C1 2.1 X 10 

-

1

2 exp (-22001T) 055 C1, 22 

73. CIO+CIO+M " CI2 +02 +M No recommendation 

- C1,23 

74a, CO+O3 

- 0100+02 1X 10"n exp(4OOO/T) +03 

74b. CIO +O0-.+OClO +O0 1 X 10 " 12 exp (40001T) 

1-1.0 

.+03 C1,24 

75 C OI 

- IO+CI -10 C1, 24 

.CI+CI0-ClO+CIO 5.9 X 10 

"'l 0A C1 

76. 

77. 

NO+OCIO 

- NO 2 + CIO 

0+0OCIO-- CIO+02 

2.5 X 10 " 1 exp (.6001T) 

2X 10 " 1 exp(-l00/T) 

0.3 

0.3 

C1,25 

C1,25 

0 

78. OH+CH3 CI CH 2CI+H 2 0 22X 10.2 exp(-l142/T) 0.1 C1,26 

79. OH + CH2 CI2 "+ CHCI2 + 

H2 0 5.2 X I0.12 exp (-10941T) 0.1 C1,26 

80. OH + CHCI 

- CC3 + H2 0 

81. OH+0CHFC 2 CFC12 +H2 0 

4.7 X 10.12 exp (-11341T) 

13X 102 exp(-l127/T) 

0.1 

0A 

C1,26 

C1,26 

82. OH+CHFCI 

- CFCI+H 

20 1.2 X 10.12 exp(-1660/T) 0.04 C1,26 

83. OH+CH2 CIF ", CHCIF +HO20 28X 10-2 exp(-1259/T) 0.1 C1,26 

84. GH+CH3 CCI 

- H*0+CH2 CC13 3.5X 10-12 exp(-1561T) 0.18 C1,26 

*Notes appear in the text. 
Reaction 

85. OH+C 

2 C14 --

C

2CI4 OH 

86. OH+C HC3 

- C0CI3 0H 

87. OH+CFCI3 "-* products 

88. OH+CCI2 -* products 


89, Br+O BrO+O 


90, O+BrO 

- Br+0 2 


91. NO+BrO 

- Br+N0 2 

92. BrO + BrO -21r + 02 

93. BrO+0 3 

- Br+20 2
 94, Br+H 2 0 2 - HBr+HO2 


95. Br+HO2 -*H~r+0 2 


96 OH+HBr HH-* 2 O+Br


97. O+H!r "* OH+Br 

98. OH+CH3 Br *. CH Br+H 

2 0 

99. CH3 +402 +M CH3 02 +M 

100. 	 CH

3

0 2 +NO --> C3 O+NO 

2 101. 	 CH3 00+ 2 - CH O+HO2 

102. 	 CH2 O+OH - HCO+H 2 0 

103. 	 CH2 0+0 

- OH+HCO 

104. 	 HCO+O 2 -* CO+HO2 

Table 1 (Continued) 

k 

9.4X 10-1 2 exp(-I199/T) 

23X 10-12 


<I X 10 

-12 exp (-3650/T) 


<1 10-12 exp (-35601T)
 F2C12 " 3.0 X 10 exp (-937/T) 

3X 	10"' 0

1 1 2.1X 

2.9 X 10-11 exp (450/T) " 4 <10 '
 

SX 10t

12 exp(-1570/T) 


5X 10"12 

2 5.1 X10' 32 7,6 X 10" exp (-1571/T) 

7.93X 10. 3 exp(_889/T) 

3.3 X !0" exp (-500/T) 1.6X 1013 exp(-3300T) 

3X 10" exp(-250/T) 

2X 10" exp(-1450/T) 

6X 10

"

12 

log k (230 K) 

0.18 

0.18 --

0.1 

0.4 

0.18 

0.18 -

03 

0.7 

0.18 

04 

0.1 

1.0 10 

0.2 

0.2 

0.2 

Note* 

CI,27 

CI,28 

CI, 29 

CI, 29 

C30 

C31 

C32 

C33 

C4 

C35,P 

C36 

C37 

C38 

C39 

DI 

D2D2 

D3 

D3 

D2 

*Notes appear in the text. 

i., 
14 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Table 2
 

Parameters for an Analytical Expression for the
 

Second-Order Rate Constant of the Reaction
 

HO + NO2 (+N2 ) -, HONO 2 (+N2 )
 

logle (k) = -AT/(B+T)-0.5 loglo (T/280) 

1 A+AZ2 + A4V
 A = A + 


B B1 + B2Z + BZ 2
 

A, = 	 31.62273 B] = -327.372 

A2 = 	 -0.258304 B2 = 44.5586 

A3 = 	-0.0889287 B3 
= -1,38092

A4 = 	 2.520173 X 10-3 

where Z -log 10 [N2 1 and is applicable only for the ranges 200 < T/K < 350 and 16.3 < " logo ([N2 I/molecule cm 3

)< 19.5, with an estimated relability in log k of± 0.10 

(reliability analogous to l o). Air is approximately 6 percent less efficient than nitrogen 

as a third body (i.e., the expression above maybe used with [Air] = 0.94 IN.]). 

NOTES TO TABLE 1 

Al 	 03 + NO -+ NO2 + 0 2 -This new recommendation is based on the 

measurements of Birks et al. (1976b), which cover the applicable 

temperature range. Earlier studies confirni the 300-K value. The 

temperature coefficient is appreciably different from that given by 

Hampson and Garvin (1975) (henceforth, noted as HG) and that 

used by the National Academy of Sciences, Panel on Atmospheric 

Chemistry (1976) (henceforth, noted as NAS). 

A2 	 0 + NO2 -*02 + NO-This recommendation is the same as that 

given by HG and NAS. There may be a slight negative temperature 

coefficient, but the evidence at low temperatures is uncertain. 

A3 	 N+O2 -+NO + O-This new recommendation is based on the measurements of Becker et al. (1969) as being the most reliable in the 

low temperature region. The value at 300 K agrees with that selected by Baulch et al- (1973b) for higher temperatures, k = 1.1 X 

10-14 T exp (-3150/T). NAS recorded this incorrectly, omitting the 

minus sign in the exponential. ­
LABORATORY MEASUREMENTS 15 

A4 	 N + NO ->N2 + 0-This new recommendation is based on the recent 

measurements of Clyne and McDermid (1975) over the temperature 

range 298 to 670 K. The uncertainty limits allow for extrapolation 

to lower temperatures. The earlier evaluation by Baulch et al. 

(1973b), cited by NAS, was based on 300-K data and high-temperature measurements on the reverse reaction. 

AS 	 OH + NO2 (+N2) -* HONO2 (+N)-This recommendation is a correlation of all recent data for this combination reaction in the 

presence of nitrogen (Anastasi and Smith, 1976; Davis, 1976; 

Anderson et al., 1974; Harris and Wayne, 1975; Howard and 

Evenson, 1974; and Atkinson et at., 1976). It emphasizes measurements at 220 K and approximately 300 K. The earlier recommendation in HG was based on data taken in the presence of other third 

bodies and corrected to nitrogen. The origin of the analytical expression cited by NAS is unknown. 

A6 	 OH + ENO3 -) H20 +NO -This new recommendation is based on 

the measurements of Smith and Zeflner (1975) over the temperature 

range 240 to 300 K. It is confirmed, above 270 K, by Margitan et al. 

(1975), whose data were cited by NAS. 

A7 	 N + NO2 -)N20 + 0-This new recommendation accepts the 298-K 

results of Clyne and MeDermid (1975). The temperature coefficient 

has been estimated here. NAS used the same 298-K value but without a temperature coefficient. 

A8 	NO + 0 + N2 -)NO 2 + N2-This new recommendation is based on 

the recent measurements of Whytock et al.(1976). It supersedes the 

earlier recommendation by Baulch et al. (1973b), accepted by HG 

and cited by NAS, in which the negative temperature coefficient is 

nearly two times larger. Other recent experiments confirm the revision (Singieton et al., 1975; and Atkinson and Pitts, 1974). 

A9 	 N+O3 -+NO+ O2-This new recommendation is based on the only 

direct measurement, which was at 300 K (Phillips and Schiff, 1962). 

The temperature dependence has been estimated here by assigning 

a pre-exponential factor similar to that for the reaction of 03 with 0 

or Cl. NAS cited the same data, without a variation with temperature, based on the older evaluation by Balch et al. (1973a). This 

new recommendaton supersedes the January 1977 estimate, k = 5 X 

10-12 exp (-650/T) cm 3 molecule-' s-1. 
16 CHLOROFLUOROMETHANE WORKSHOP REPORT 

A10 	 NO 2 + 03 -> NO3 + 0 2 -This recommendation is the same as that 

used by NAS and recommended by HG. It is based on an equally 

weighted least-squares fit to the 22 points in the studies of Davis 

et al. (1974), Graham and Johnston (1974), and Huie and Herron 

(1974), all of which include measurements below 273 K. The older 

evaluation by Baulch et a!. (1973b) shows a much larger temperature 

coefficient, based on old data. 

All 	 HO2 + NO 2 + N2 -?. H0 2N0 +N -Peroxynitric acid, HOON0 2, 

is a potential sink for NOx (and 110) in the atmosphere. Its rate of 

formation has been reported by Simonaitis and Heicklen (1977) and 

Howard (1977). Niki et al. (1977) have reported the infrared spectrum of the product of the above reaction. 

A tentativerecommendationis made here for this formation reaction. 

However, peroxynitric acid should not be introduced into atmospheric models until there is also kinetic information on its photolysis 

and reactions with atoms and free radicals; it may be an ineffective 

sink. 

The tentative recommendation is that the rate constant to be used 

for HO2 + NO -> HO2NO 2 be the expression developed by 

Zahniser et al. 6977) for CIO + NO 2 CIONO 2 , with twice the 

uncertainty assigned to that reaction. 

Al2 	 Reactions of 0 (1D)-The recommendations adopt the time-resolved 

0 ('D) emission measurements at the National Oceanic and Atmospheric Administration (NOAA) Laboratories, Boulder, Colorado, for 

reactions with 02' 03, N2 , N20, NH, H20, CO2 CH and HC1 

(Streit et al., 1976; and Howard, private communication, 1976), 

which supersede earlier reports. The branching ratio for the reaction 

with N20 is based on numerous earlier measurements. See 

Cvetanovic (1974) for a review. However, the ratio k12/k 1 3 may be 

slightly less than unity, based on earlier results of Simonaitis et al. 

(1972) and recent, unpublished results from the NOAA laboratories. 

The rate constant for 0 + NZ + M -- N 0 + M is taken from the 

recent study by Kajimoto and Cvwtanovic 6976). 

The reactions with CFCl3 and CF2C12 have been studied only by the 

resonance absorption technique at 300 K, as part of a series of 

studies at the University of Cambridge (Fletcher and Husain, 1976a, 

b; Heidner and Husain, 1973; and Heidner et al., 1973). When these 

absorption studies overlap those of the NOAA Laboratories, there is 
LABORATORY MEASUREMENTS 17 

usually a systematic difference. The reported results have been 

scaled by multiplying them by 0A1, and a zero temperature 

dependence has been assigned by analogy to the other chemical 

reactions of 0 ('D). The chlorocarbon rate constants are for total 

disappearance of 0 (1D) and probably include physical quenching. 

The lower limit for the fraction in the reactive channel forming ClO 

is 039 (CFCI3) and 0.47 (CF 2 CI2 ) (Gillespie et al., 1977). 

For CCI20, CFCIO, and CF10, the only reported values are relative 

to the value for 0 CD) + N2 0. The recommended values combine 

these relative values by Jayanty et al. (1976) with the value given 

here for O (1D)+ N2O (1.1 X 10-10). However, these values may be 

low, since earlier measurements by the same group with the same 

technique gave ratios for CFC13 and CFCl2 relative to N.0 that 

were a factor of 2 lower than the same ratios measured by Pitts et al. 

(1974) and also derived from the absolute values of Fletcher and 

Husain (19 7 6 a) 

It must be emphasized that there exist two disparate sets of absolute 

rate data for O(1D)reactions. Although we have chosen to accept as 

correct the results obtained by the 0 (1D) -§ 0 (

3P) emission measurements in the NOAA Laboratories, the possibility cannot be ruled 

out that the results of the resonance absorption studies at the University of Cambridge are correct. 

B1 	 In the January 1977 version of this report, we attempted to resolve 

the conflicting data obtained from ratio and direct measurements on 

reactions 29 through 31 (O + H0 2 -H 2 0+0 (29), HO 2 + H02 

-+ H202 + 02 (30), and H0. + NO --OH + N02 (31)) and to formulate a consistent set of rate constants. That procedure gave a great 

deal of weight to: (a) studies of k29 and k3, that did not involve k3, 

(and obtained "low" rate constants), and (b) ratio determinations of 

k 9 k30112 and k31/k30 1 2 . It was decided that the existing data 

could be brought into agreement if a lower value of k were 

1
 adopted-i X 1012 instead of 3.6 X 10- 12 cm3 s- .Subsequent to 

that evaluation, Howard (1977) has measured k31 directly in his laser 

magnetic resonance system and obtained a result nearly 30 times 

"
 higher than the earlier recommendation (8 X 1012 versus 3.5 X 10-13 

cm3 s-l). Because this measurement represents the best information 

available for any of these three reactions, it must form the nucleus 

for the new recommendations. The value recommended here is 

temperature-independent, consistent with its large magnitude and 
18 CHLOROFLUOROMETHANE WORKSHOP REPORT 

the determination of Glanzer and Troe (1975) near 1500 K (7.5 X 

10-12 cm 3 S-'). 

The basis for the previous recommendation for k.1 was primarily the 

study of Hack et at. (1975). In light of Howard's measurement, 

however, it appears that their study may be in error due to reaction 

31 itself. They used the H + NO2 -* OH + NO reaction to form OH - at concentrations of"Iv 3 X 1013 Cem3 and subsequently formed HO2 

via OH + H2 0 2 -> HO2 + H2 0. The higher value for k31 makes it 

very likely that HO2 was being reconverted to OH via reaction 31, 

using the NO formed in the OH formation step. Because the decay 

of OH was followed for less than a factor of 2 change,it is difficult 

to assess to what extent this back reaction decreased the apparent 

rate constant. Other indirect ratio measurements for k31 also need to 

be reevaluated to ascertain the extent to which their earlier assumptions are still realistic. For example, the new value for k31 is not 

compatible with the Simonaitis and Heicklen (1976a) determination 

of k31/k3 1/2 because a value of k3 = 3 X 10-10 would be required. 

Clearly, their study, as well as others, needs to be reexamined in light 

of the drastic change in k31 to determine if any of their kinetic assumptions have been invalidated. 

Because the ratio determinations involving the HO2 + NO reaction - are in question, there is no basis for the low value of k30 (1 X 10 2 

cm3 s-1). The recommendation is to now use the k/a ratio of 

Hamilton and Lii (1976) and the u measurement of Paukert and 

3 s- Johnston (1972). The resulting k 3 = 2.5 X 10-12 cm I is somewhat lower than the Paukert and Johnston value and may indicate 

that their a is also high by '\ 30 percent, although there is insufficient 

data to draw any firm conclusions. In view of the rather poor state 

of knowledge for this reaction, no temperature dependence is recommended. 

The recommendation for k29 is unchanged forn the earlier report 

(3 X 10-" cm3 s0). ­ This value is consistent with the upper limits of 

Chang and Kaufman (1976) and Howard (1977) and is within our 

estimate of the experimental uncertainty of the recent Burrows et al. 

(1977) result of 5.1 X 10-11 em 3 S-. Clearly, additional unambiguous studies are sorely needed on all three reactions. 

B2 Both Simonaitis and Heicklen (1973) and DeMore and TschuikowRoux (1974) measured the ratio Ik(HO2 + O3) 1/2. [k (H02 + HO2)] 
LABORATORY MEASUREMENTS 	 19 

The average value for this ratio is 4-6 X 10-8 exp (-1275/T) cm 1 /

2 

molecule- 112 . Using this ratio and the value of k(HO2 + HO2 ) recommended in this evaluation (see note Bi), the value for k(HO2 + 

03) given in table 1 was obtained. This value must be changed if 

there are any changes in k (HO2 + HO2 ). The tentative January 

1977 value for HO2 + O3, based on the earlier k(HO2 + HO2 ), was 

I X 10- 3 exp (-1525/T) cm3 molecule- s2. 

B3 	 The A-factor was decreased by 8 percent from that in Hampson and 

Garvin (1975) to correct 03 absorption measurements to standard 

atmospheres. The activation energy is that of Anderson and 

Kaufman (1973), which is 500 caloriesmole lower than the more 

recent DeMore (1975) work, but is preferred because of the larger 

temperature range (220 to 450 K) covered by them versus the more 

limited range (270 to 335 K) in the DeMore work. 

B4 	 There are no data on the temperature dependence of either reaction. 

The recommedded value for 0 + OH is the same as that of NBS 

TN866 since no new measurements are available. The recommended 

value for 0 + HO2 is based on the recent measurements of Burrows, 

etal.(1977). TheJanuary 1977values for both 0 + OH and 0 + HO 2 

were 	 - 1 X 10-10 exp (-250/T) cm 3 molecule ' s-1. 

B5 	 Although there is no evidence, products are most likely OH + HO2 , 

because the H2 0+02 requires a complex rearrangement. More work 

is needed, especially below room temperature. The present recommendation is the same as that of NBS TN 866. 

B6 	 This reaction is one of the few that does not require further work. 

New results ,are in excellent agreement with previously established 

values. The recommendation is the same as that of NBS TN 866. 

B7 	 More work is needed, especially below room temperature. The value 

here is a report by Clyne and Monkhouse (197-7) measured over the 

300- to 600-K range. 

B8 	 More work is needed, especially on the T dependence around and 

below room temperature. The recommendation is the same as that 

of NBS TN 866. 

B9 	 The value employed by NAS of B = 2500 K (B = E/R) is very large 

for arecombination involving four atoms. For comparison, OH + NO 

and OH + NO 2 have B -- 900 K. In addition, k (295) = 1.2 X 10-29 
20 CHLOROFLUOROMETHANE WORKSHOP REPORT 

B1O 

B11 


C! 


C2 

C3 

cm 6 s-1 from the NBS expression, which is much larger than 2.5 X 

cm 6 ­ 10-31 s1 from Trainor and von Rosenburg(l974). The expression here fits their k (295) and uses B = 900 K. Clearly, more work 

is needed, especially on temperature dependence. 

More work is needed, especially below room temperature. This value 

is a composite 670, of a recent Hack et al. (1975) measurement of 8 X

10-12 exp (---0 cm3 molecule-1 0 with earlier work of Greiner 

(1968). Although the two studies are in relatively good agreement, 

there are reasons to question both determinations. The Greiner work 

involved a large temperature increase due to absorption of flash 

3
 energy. TheHacket al. study used radical densities of 3 X 1013 cm

and may have been complicated by the back reaction between the 

product HO2 andresidualNO from the OH formation step. The new 

value for k (HO2 + NO) of 8 X 10- 12 implies a very rapid reconversion. Additional studies are needed. 

Several recent investigations of reaction 42 at atmospheric pressure 

in noninert gases have found an enhancement of k of about a factor 

of 2 (Sie et al, 1976; Cox et a., 1976; Chan et a., 1977; and 

Atkinson et al., 1977). There is virtually no pressure dependence 

with He or Ar. For stratospheric calculations, the pressure dependence may be disregarded; however, in tropospheric models it must be 

taken into account. Ina recent paper, Smith (1977) explains the 

pressure effect, temperature independence, and low absolute magnitude as consistent with an energized HOCO radical as an intermediate. 

Clearly, more experiments are needed utilizing various M gases and 

both reactant and product identification. 

All of the data used in this evaluation are tabulated in the reviews of 

R. T. Watson (1974, 1977) and the position paper entitled "Review 

of Halogen Kinetics." 

This evaluation was based on data obtained from Watson et al. 

(1976), Zahniser et al. (1976), Kurylo and Braun (1976), and Clyne 

and-Nip (197 6 a). 

-1 1 This expression is based on values of 5D X 1011 cm 3 molecule s1011 "
 and 4.4 X cm3 molecule-' s- at 298 and 230 K, respectively. These values were deduced from the experimental data 

of Bemand et al. (1973), Clyne and Nip '(1976b), and Zahniser 

and Kaufman (1977). 
LABORATORY MEASUREMENTS 21 

C4 	 The recommended rate constant is an evaluated expression that has 

a 298-K value equal to the average of two available experimental 

measurements (Clyne and Watson, 1974; and Zahniser and Kaufman, 

1977), and that has a temperature dependence as measured by 

Zahniser and Kaufman (1977). The tentative January 1977 value 

was 2.2 X 10-11 Cm3 molecule- 1 s-1 . 

C5 	 This evaluation was based on the data of Anderson et al. (1974), 

Smith and Zellner (1974), and Ravishankara et al. (1977). 

C6 	 This value is based on the four recent direct studies that measured k 

at stratospheric temperature (Watson et al., 1976; Manning and 

Kurylo, 1977; Zahniser and Kaufman, 1977; and Whytock et al., 

1977a). It does not consider the discharge flow-mass spectrometric 

results that determined k at and above 300 K (Clyne and Walker, 

1973; 	Poulet et al., 1974; and DeMote and Leu*), and the competitive chlorination results that measured B (CH 4 ) - B (C2H6) from 190 

to 650 K (a recent study by DeMore and Lint agrees with the literature value (Fettis and Knox, 1964) of 1400 K for B (CH 4 ) - B 

(C'2H6)). It was recently reported by both Whytock et al. (1977a) 

and Zahniser and Kaufman (1977) that the Cl + CH4 reaction exhibited a nonlinear Arrhenius behavior, which appeared to explain 

the differences in B (CH4) between the studies that determined B 

below 300 K and those above 300 K (even though it is acknowledged 

that the absolute values of k above 300 K in the mass-spectrometic 

studies and the resonance fluorescence studies are not in particularly 

good agreement). However, the competitive chlorination results 

indicate that the magnitude of B (CH 4 ) - B (C2116) is constant with 

temperature, which strongly indicates that the Arrhenius behavior of 

C1 + CH 4 is linear (the Cl + CH 4 and Cl + C2H6 reactions are unlikely 

to have identical nonlinear Arrhenius behaviors). The value of B for 

C2H6 has been reported to be 60 K, yielding a value of 1460 K for 

B (CE 4). However, if the resonance fluorescence results have underestimated B (CH 4 ), it is possible that the resonance fluorescence 

results for B (C2 H6 ) have been underestimated as well. In conclusion, it can be stated that no significant errors can be found in any 

particular study; therefore, the direct data collected at stratospheric 

temperatures will be used until further data are available. Note the 

unsymmetric error limits that allow for a wide range of k48 at 230 K. 

*W. B. DeMore and M. T. Leu, let Propulsion Laboratory, Pasadena, California, 

unpublished data, 1977. 

tW. B. 	 DeMore and C. L. Lin, Jet Propulsion Laboratory, Pasadena, California, 1977. 
22 CHLOROFLUOROMETHANE WORKSHOP REPORT 

C7 	 These data were reported by Leu and DeMore (1976) 

C8 	 Thisis an evaluated expression put forward by Zahniser et al. (1977). 

It is consistent with the available low-pressure data (Leu et al., 1977; 

Zahniser and Kaufman, 1977; and Birks*). The pressure dependence 

is based on a theoretical RRKM model of Smith and Golden (1977). 

No experimental measurements of falloff behavior have been reported. The tentative January 1977 value was 5.1 X 10-33 exp 

6 2 1 (+10331T) cm molecule " s , with no pressure falloff 

C9 	 This is a composite expression derived by Kurylo (1977) based on 

his data combined with that reported by Molina et al. (1977). The - January 	 1977 recommendation was 4.5 X 10 12 exp (-840/T) cm3 

molecule-1 -1 s . 

CIO 	 This value was reported by Kaufman and Zahniser.t 

Cll 	 This value is based on the data reported by Kurylo arid Manning.4 

C12 	 This evaluation is based on the results of Wong and Belles (1972), 

Brown and Smith (1975), Ravishankara et at. (1977), and Hack et al. 

(1976). The results of Balakhnin et al. (1971) have not been used. 

C13 	 This value combined the evaluation of k. with the most recent 

thermodynamic data for AHf and So of HCI, 0 OlII), f and CI(2P). 

C14 	 This value is based on the results obtained below 300 Kin the recent 

studies reported by Watson et alt (1977) and Whytock et al. (1977b). 

However, if the data of Benson et al. (1969) and Westenberg and 

deHaas (1968) are included with all the data of Watson et al. (1977) 

and Whytock et al. (1977b), the overall expression is almost identical. 

C15 	 This expression is based on the activation energy reported by 

Michael et al. (1977) and an A-value that has been modified to yield 

the mean of the values reported for k57 at 298 K by Michael et al. 

(1977) and Leu and DeMore (1976). This value is in good agreement 

with that reported by Watson et al. (1976). Two reaction channels 

*J. Birks, Univertity of 1uinois, unpublished data, 1977. 

tF. Kaufman and M.S. Zahniser, University of Pittsburgh, unpnblished results, 1977. 

tM- J. Karylo and R. Manning, National Bureau of Standards, Washington, D.C., 

published data, 1977. 
LABORATORY MEASUREMENTS 23 

are exothermic; however, there is no experimental evidence for 

process (b). 

C16 	 This is the estimated pre-exponential A-value to yield the 298-K 

value reported by Leu and DeMore (1976). 

C17 	 The reliability of the published values of k62 are certainly questionable, and there is no experimental data for k6. However, the photochemical models require only Kk= k6k6, and not individual values 

for k62 and k63- t ough can be calculated from thermochemical data, the best literature value for AH C100 is uncertain by 

1 ±5 kJ 	mol " , leading to an uncertainty of a factor of 10 in the value 

of Keq 

CIS 	 The rate constants are not well-defined for either channel. The value 

shown for channel (a) is that reported by Johnston et al. (1969), 

whereas the value for (b) is approximately an order of magnitude 

greater than that reported by Johnston et al. and is influenced by 

two factors: (1) the ratio of k64a/k64b reported by Nicholas and 

Norrish (1968), and (2) the rate constant for k7,, and thermochemical data, assuming that E3(64b) is zero. 

C19 	 There is no recommendation because there is no experimental data. 

C20 	 There is no experimental data. k has been assumed to be intermediate between k (l0 2 + HO2) and k (C10 + CIO). 

C21 	 These are estimated Arrhenius expressions based on the data reported 

by Clyne and Watson (1974). 

C22 	 Although the branching ratio between reactions 72a and 72b is not 

well-defined, the overall rate constant (72a + 72b) is accurately 

known at low pressures. It is assumed in this review that the branching ratio is 2:1. There is experimental evidence for reaction 72c, but 

this expression was derived by combining thermodynamic data with 

the value of k7 ,. The low precision is due to the uncertainty associated with the value of AHf 298 for OCO. 

C23 	 The experimental data base does not strongly support the presence 

of a third-order reaction channel for CIO' decay. 
24 CHLOROFLUOROMETHANE WORKSHOP REPORT 

C24 	 The branching ratio between channels (a) and (b) is not well-defined 

but, for the present discussion,is assumed to be unity. The Arrhenius 

expressions were estimated on the basis of data reported by DeMore 

and Lin* at 298 K. 

C25 	 These are estimated Arrhenius expressions based on 298-K data. 

C26 	 These are evaluated expressions based on all available data, which are 

normally in good agreement. 

C27 	 This value was reported by Chang and Kaufman (1977). 

C28 	 Although all three studies of k86 are in excellent agreement at 298 K, 

there is some doubt as to the temperature dependence. However, 

the uncertainty limits adequately allow for any temperature dependence within the 200- to 300-Krange. 

C29 	 These are the estimated upper limit Arrhenius expressions based on 

the 480-K data of Kaufman and Chang.t 

C30 	 This is the Arrhenius expression of DeMore and Leu.t There is excellent agreement between this work and that of Clyne and Watson 

(1975) at 298 K. 

C31 	 This value is based on data of Clyne et al. (1976). 

C32 	 This value is based on data of Clyne and Watson (1975). 

C33 	 This is the activation energy reported by Clyne and Cruse (1970), 

but the A-value has been modified to yield the value of k92 (298 K) 

reported by Clyne and Watson (1975). 

C34 	 This is the upper limit set by Clyne and Cruse (1970); there is no 

evidence for reaction. 

C35 	 This is the estimated Arrhenius expression based on the value of k94 

(298 K) reported by DeMore and Leu.4 

*W. B. DeMore and C. L. Lin, Jet Propulsion Laboratory, Pasadena, California, 

unpublished data, 1977. 

tF. Kaufman and J. S. Chang, University of Pittsburgh, unpublished data, 1977. 

$W. B. DeMore and M. T. Leu, let Propulsion Laboratory, Pasadena, California, 

unpublished data, 1977. 
LABORATORY MEASUREMENTS 25 

C36 This is an estimate; there are no experimental data.
 

C37 This value is based on the data of Takacs and Glass (1973a, b).
 

C38 This evaluation is based on the data of Brown and Smith-(1975) and
 

Takacs and Glass (1973c). 

C39 This value is based on the data of Davis et al. (1976b). 

DI This is the sole fate of CH 3 radicals in the atmosphere; reaction 

proceeds at the rate of reaction 43. 

D2 See Hampson and Garvin (1975). 

D3 This is the A-factor chosen and activation energy derived to fit roomtemperature value in Hampson and Garvin (1975). 

CROSS SECTIONS AND QUANTUM YIELDS 

INTRODUCTION 

This subsection is concerned with photoabsorption cross sections and with 

quantum yields for photodissociation in the visible and ultraviolet wavelength regions for molecules that are of importance in the stratosphere. 

Several excellent reviews on this subject have been published recently: 

* 	 "Absorption Cross Sections of Stratospheric Molecules," by Hudson 

and Kieffer (1975); Section 5.8 of the Climatic Impact Assessment 

Program (CIAP) Monograph I 

* 	 "Rate and Photochemical Data," by Hampson and Garvin (1975); 

Section 5.9 of CIAP Monograph I, which contains an extensive table 

providing information on quantum yields and' absorption cross 

sections 

* 	 "Photodissociation Ratesin the Atmosphere Below 100 Kilometers," 

by Turco (1975); includes a discussion on absorption cross sections 

and quantum yields 

* 	 "Rate Constants of CIOg of Atmospheric Interest," by Watson 

(1977); reviews absorption cross sections and quantum yields for 

molecules containing chlorine atoms 
26 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Table 3 lists photochemical reactions of stratospheric interest, divided into 

two groups: (1) reactions referred to the foregoing reviews, and (2) reactions for which there are additional data not included in the earlier reviews. 

Only reactions pertaining to the second group are discussed here; reference 

is made only to those aspects of the reactions for which new or supplementary information is available. Details of measurements, analysis of 

errors, etc., are not repeated here; see Hudson (1971). 

The absorption cross sections reported here are defmed by the following 

expression of Beer's Law: 

I = Io exp ( uc2) 

where 

I,1 	 = incident and transmitted light intensity = o absorption cross section, cm2 molecule-1 

c = concentration, molecule cm-3 

£ = pathlength, cm 

DISCUSSION OF PHOTODISSOCIATION REACTIONS 

Species Containing N, H, and 0 Atoms 

03 + h -+ 0 (

TD) + 0 2 (Reaction3) 

The photolysis of ozone at wavelengths near 300 nanometers to produce 

electronically excited oxygen atoms, 0(1D), is a very important process in 

the atmosphere; the reaction of O('D) with water is the principal source of 

hydroxyl radicals; and nitric oxide is produced in the stratosphere by the 

reaction of O(D) with N2 0. 

Several groups have reinvestigated the photo dissociation of ozone since it 

was reviewed by Hampson (1973); table 4 (taken from Philen et al., 1977) 

summarizes the results of quantum yield measurements, 0 (Ot1)), at 313 

nanometers. As table 4 shows, the disagreement is very significant: the 

reported (O1D) values range from 0.1 to 1.0 at 298 K. The studies that 

report 0 (01D) as a function of wavelength-do not resolve these discrepancies. 

The tentative recommendation for quantum yields at 235 K is given in 

table 5; this recommendation is based on the earlier work of Liu and 
LABORATORY MEASUREMENTS 27 

Table 3 

Photochemical Reactions of Stratospheric Interest 

1. 02 +bp-*O + 0 (a) 25. C10 3 + h-- products 

2. 03 + hv- O + 0 2 (a) 26. HCI+ h - H + C1 (c) 

3. 03 + hv-* O(1D) + 0 2 27. HOCI + h --OH + CI 

4. NO +l + N + 0 (a) 28. CINO + hv- CI + NO 

5. NO2 + hw- NO + 0 29. CINO 2 1w-> products 

6. NO3 + hv-- NO + 02 30. CIONO + hv+ products 

7. NO3 +hv - NO 2 +0 31. CIONO 2 + hz - products 

8. N2 0 +hv-+N 2 + O('D) 32. Cl2 + hvP- CI + C1 (c) 

9, N20 + hu-- products 33, C12 0 + hv-4 CI + CIO (c) 

10. NH 3 + hP ' NH 2 + H (a) 34. CCI4 + hv -+products (c) 

11. HO 2 + hP - products (b) 35. CC13F + h -* products 

12. H20 + hP H + OH (a) 36 CCI2IF +1h -'products 2 

13. H202 + hy * OH + OH (a) 37. CCIF3 + hv -* products 

14. HNO2 + hv "+OH + NO 38. CHCI2F + hP -* products (c) 

15. HNO 3 + hv --OH + NO 2 (a) 39, CHCIF 2 + hv --products (c) 

16. SO2 + lw-* SO + 0 (a) 40. CH 2 CIF + hP - products (c) 

\17. H2S + hv- HS + H (b) 41. CHsCI + hi, --'products (c) 

18. CO + hv - C + 0 (a) 42. CCI2 FCCIF2 + hv - products 

19. CO2 + h, - CO + 0 (a) 43. CC1FPCCIF 2 + h -* products 

20. CH4 + hv - products (b) 44. CCIF CF3 + hv - products 

21. CH 2 0 + hi -0 products (a) 45. CH3CCI3 + hv - products 

22. CIO +hv-Cl +O 46. CCIO + hu-- products 

23. CIO 2 + Iw -+products (c) 47. CCIFO + hv -* products 

24. OCIO + h - 0 + CIO (c) 48. CF20 + h--- products 

(a) Hudson and Kieffer (1975) and Hampson and Garvin (1975). 

(b) Turco(1975). 

(c)Watson (1977). 
Table 4
 

Summary of Reported Quantum Yields for 0(1 D) Production
 

from Photolysis of Ozone at 313 Nanometers 

Investigator 4)O('0) (K)emperature_____

Castellano and Schumacher 


(1969; 1975) 


Kajimoto and Cvetanovic 


(1976) 


Simonaitis et al. (1973) 


Martin et al. (1974) 


Moortgat and Warneck 


(1975)


Kms at al. (1975) 


Jones and Wayne (1970) 

Lm and DeMote (1973174) 

Philen et al, (1977) 

Temperature (K) 

0.75 248 

10 298 

0.53 313 

0.21 198 

0.35 298 

032 298 

0 29 298 

0.29 293 

0.22 258 

0.11 221 

0.1 298 

008 235 

0.12 298 

Technique 

Gas phase 0.; decrease 

in 03 pressure 

Gas phase 03, N2 0, 

increase in N2 

Gas phase 03, N2 0,0 

increase in N2 

03 N20; NO 2 chemiluminescence 

03, N2 0; NO2 chemiluminescence
 

Gas phase 0., 02, N20; 


increase in N2 

Decrease in 03 

03, isobutane 

03 ,N 2 0;N chemilumi- 2 

nescence 

Photolytic Source* 

CW lamp 

0 

W 

CW lamp; 

chemical filter 

CW lamp; 

chemical filter 

0 

0 

CW lamp 

CW lamp 0 

CW lamp; 0 

chemical filter 

CW lamp 

CW; monochromator 

Pulsed dye 

laser 

*CW = continuous irradiation (usually with medium pressure Hg lamp). 
Table 5
 

Quantum Yields for 0(D) Production
 

from Photolysis of Ozone at 235 K
 

(rm) 00('D) * (nm) 

<303 1.00 310

304 0.99 311 

305 0.95 312 

306 0.90 113 

307 0.80 314 

308 0.55 315 

309 0.35 > 316 

*Relative to 0 = 1 for X< 300 nanometers. 

0 (1D) * 

0.25 0 

0.15 

0.10 

0.07 

0.04 

0.02 

0 

to 
30 CHLOROFLUOROMETHANE WORKSHOP REPORT 

DeMore (1973), who measured quantum yields as a function of wavelength 

at 235 K, using a light source with a bandwidth of 1.6 nanometers. If the 

temperature effect is taken into account, their results are in good agreement with those of Philen et al. (1977), who determined quantum yields 

at 298 K as a function of wavelength, with a photolytic source having a 

much narrower bandwidth (about 0.1 nanometer). 

The values shown in table 5 were estimated from the figure given by Lin 

and DeMote (1973). This recommendation is rather arbitrary; no explanation can be forwarded for the significant discrepancies at 313 nanometers 

between the value of 0.07 given in table 5 and the values reported by Kuis 

et al. (1975), Kajimoto and Cvetanovic (1976), Martin et al. (1974), and 

Moortgat and Warneck (1975), which are a factor of 2 to 3 larger. Further 

work is probably needed to resolve this issue. 

The quantum yields noted in table 5 were based on the assumption that 

the quantum yield below 300 nanometers is unity. However, a recent 

study by Lawrence and Stone (private communication, 1976) has indicated 

that 0 (013)has a value between 0.87 and 0.93 for the wavelength region 

274 to 300 nanometers. Therefore, if 0 (0'D) is less than unity at 300 

nanometers, the relative measurements above 300 nanometers must be 

modified accordingly. 

NO2 + hp -+ NO + 0 (Reaction5) 

The absorption cross sections for nitrogen dioxide have been remeasured 

as a function of temperature by Bass et al- (1976), who report extinction 

coefficients every 1/8th nanometer between 185 and 410 nanometers at 

298 K, and between 290 and 400 nanometers at 235 K. The effect of the 

dimer (N2 0 4 ) absorption was considered in detail, and the measurements are probably correct to within ± 10 percent. Table 6 lists these 

results for every 5 nanometers. 

NO3 +hv -+ NO + 02 (Reaction 6) and NO 3 +h --+NO+ 0 (Reaction7) 

The photoabsorption cross sections of the nitrogen-trioxide radical, NO3 , 

taken from Graham (1975) are included as table 7; these results supersede 

the earlier ones of Johnston and Graham (1974). The quantum yields, 

which are wavelength dependent for each of the channels (i.e., for reactions 

6 and 7), are rather uncertain. For atmospheric calculations, it is recommended to use simply the photodissociation'rates estimated by Graham - (1975); for reaction 6,J 0.04 see t ; and for reaction 7, J 0.10 sec t 

(at the Earth's surface, for overhead Sun). 
LABORATORY MEASUREMENTS 31 

Table 6
 

NO2 Absorption Cross Sections at 235 and 298 K
 

a (cm2) a (cm2

) 

(nm) 235 K 298 K (nm) 235 K 298 K 

185 260(-19) 300 1 09(-19) 1.17(-19) 

190 293 305 1 67 1 66 

195 2A2 310 1.83 1.76 

200 250 315 2.19 2.25 

205 3.75 320 2.35 2.54 

210 385 325 2.54 279 

215 402 330 291 2.99 

220 3.96 335 3.14 3.45 

225 3.24 340 323 3.88 

230 243 345 3.43 407 

235 1.48 350 3 11 4.10 

240 6.70 355 437 5.13 

245 4.35 (-20) 360 390 451 

250 2.83 365 5.37 5 78 

255 145 370 487 542 

260 1.90 375 5.00 535 

265 201 380 593 599 

270 3 13 385 5 79 594 

275 402 390 549 600 

280 5 54 395 562 5.89 

285 6.99 400 666 676 

290 677(-20) 818 405 5-96 632 

295 8.52 967 410 532 5.77 

N2 0 +h N2 + 0 (Reaction 8) 

The ultraviolet absorption spectrum of N20 has been reinvestigated by 

Johnston and Selwyn (1975); their results agree with the previously measured ones in the range 210 to 235 nanometers (Hudson and Kieffer, 1975), 

but disagree with Bates and Hays (1967) for wavelengths above 260 nanometers. The significant new result of Johnston and Selwyn (1975) is that 

the cross sections at these longer wavelengths are vanishingly small, a result 
32 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Table 7
 

NO3 Absorption Cross Sections from 400 to 700 Nanometers*
 

w 10t9 0 x. 10190 A 10190 -A l.o1ll u A* loll 

400 00 460 3.9 520 144 580 29.9 640 9 8 

402 0.1 462 3.5 522 17.2 582 310 642 68 

404 0.2 464 4.1 524 15.0 584 24.7 644 7.1 

406 0.3 466 4.5 526 13.7 586 275 646 5.6 

408 0.3 468 50 528 179 588 448 648 4.3 

410 0.6 470 4.9 530 209 590 56.7 650 3.2 

412 03 472 5.4 532 18.1 592 483 652 3.9 

414 0.7 474 5.6 534 17.7 594 39.2 654 5.7 

416 0.3 476 6.4 536 23.2 596 416 656 89 

418 0.6 478 66 538 211 598 35.4 658 16.8 

420 0.9 480 6.4 540 18.1 600 24.5 660 51 2 

422 0.8 482 6.3 542 168 602 28.4 662 1708 

424 1.2 484 6.2 544 13.9 604 40.0 664 115.4 

426 0.9 486 7.4 546 20.4 606 33.8 666 486 

428 1.2 488 80 548 27.5 608 15.9 668 17.5 

430 1 2 490 93 550 22.4 610 13 5 670 75 

432 14 492 8.9 552 21.6 612 16.9 672 5.7 

434 I 7 494 8.8 554 24.5 614 22.4 674 3.6 

436 2.1 496 104 556 29.5 616 17.4 676 3.1 

438 1.8 498 10.8 558 31.7 618 18.3 678 5.5 

440 1.9 500 9.8 560 32.3 620 24.7 680 4.9 

442 20 502 9.1 562 268 622 761 682 2.5 

444 2.1 504 10.5 564 24.8 624 116.6 684 0.9 

446 2.3 506 11.9 566 25.8 626 70.0 686 0.3 

448 2.8 508 10.6 568 25.7 628 68.9 688 0.4 

450 2.7 510 130 570 25.3 630 64.1 690 0.1 

452 3.1 512 16 1 572 24,8 632 32.7 692 0.0 

454 34 514 14.1 574 25.5 634 13.2 694 01 

456 3.2 516 14.0 576 29.2 636 123 696 0.4 

458 3.7 518 12.1 578 303 638 17.6 698 0.A 

*Xin nm, u in cm 2 . 
LABORATORY MEASUREMENTS 33 

that implies a negligible rate of photolysis of N20 in the troposphere. 

This result has been corroborated by Stedman et al. (1976), who measured 

directly an upper limit to the N20 photodissociation rate at ground level - of 10 sec7. 

N2 0 5 + hv -+ Products (Reaction 9) 

Table 8 lists data for N2 05 taken from Graham (1975), which supersedes 

the results from the review article by Johnston and Graham (1974). The 

quantum yields for photodissociation are unknown; possible products are 

NO 2 +NO3 and N2 0 4 +0. 

HNO2 +hp -* Products(Reaction 14) 

Absorption cross sections of UNO have been measured in the region of 

200 to 400 nanometers by Graham e8975) and Cox et al. (1976). Graham's 

original measurements gave relative numbers that should be multiplied by 

a factor of 2.4 (Graham, private communication, 1976) to bring them into 

agreement with the measurement of Zafonte (1975), which provides an 

HNO2 -photolysis rate relative to an NO2 photolysis rate under conditions 

of broadband excitation (310 to 410 nanomet&s). The cross sections 

obtained in thisway are at least a factor of 2 smaller than the data of Cox 

et al. (1976); no preferred values can be given. The two sets of data are 

listed in table 9. 

Chlorine Containing Species 

ClO -/h - CI + 0 (Reaction 22) 

The absorption cross sections of chlorine monoxide, CIO, have boen-reviewed by Watson (1974). Two recent calculations (Langhoffet al, 1976; 

and Coxon et al., 1976) indicate that photodecomposition (predissociation 

of 'the A2 1-1 2 state) of ClO accounts for at most 2 to 3 percent of the 

total destruction rate of CIO, in the stratosphere, which occurs predominantly by reaction with oxygen atoms and nitric oxide. 

C/o3 +hv -* Products (Reaction 26)' 

Table 10 lists absorption cross sections of chlorine trioxide, C10 3 , for the 

200- to 350-nanometer range obtained by graphical interpolation between 

the data points of Goodeve and Richardson (1937). Although the quantum 
34 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Table 8 

N2O 5 Absorption Cross Sections from 200 to 400 Nanometers 

x a x a xa 

(nm) (cm2) (in) (cm 2) (nm) (cm2) 

206 6.6 (-18) 246 43 (-19) 286 7.8 (-20) 

208 5.9 248 3.8 288 7.1 

210 5.2 250 3.5 290 6.3 

212 4A 252 3.0 292 5.7 

214 3.7 254 2.72 294 4.9 

216 3.0 256 2.55 296 4A 

218 2.48 258 2.33 298 3.8 

220 2.06 260 2.12 300 3.2 

222 1.71 262 1.97 302 2.7 

224 1.41 264 1.86 304 2.4 

226 1.23 266 1.7 306 2.1 

228 1.06 268 1.64 308 1.8 

230 9.3(-19) 270 1.52 310 1.5 

232 8.4 272 1,42 320 7.5 (-21) 

234 7,5 274 1.31 330 4.0 

236 6.9 276 1.2 340 2.7 

238 6.3 278 1.15 350 1.8 

240 5.7 280 1.07 360 1.0 

242 53 282 9.9 (-20) 370 4.7 (-22) 

244 4.7 284 8.9 380. 1.3 
LABORATORY MEASUREMENTS 35 

Table 9 

HNO 2 Absorption Cross Sections from 200 to 400 Nanometers 

1. 200- to 300-namometer regi o n (a) 

X 102 X I02 1020 u 

(nm) (cm 2) (nm) (Cm2) (nrm) (Cm2 ) 

200 125 235 68 270 2.0 

205 185 240 42 275 < 0.4 

210 295 245 28- 280 < 0.4 

215 425 250 18 285 < 0.4 

220 177 255 11 290 <04 

225 140 260 8 295 < 0.4 

230 104 265 4 300 <04 

2. 300- to 400-namometer region 

100 g (0m2) 1020 a (e&) 

(nm) (a) (b) (nm) (a) (b) 

300 < 0.4 1.63 352 28 10.2 

302 1.42 354 56 24.2 

304 1.68 356 20 11.7 

306 2.09 358 19 7.73 

308 1.99 360 17 6.10 

310 <0.4 2.18 362 16 5.21 

312 < 0.4 2.64 364 21 8.09 

314 2.0 3.00 366 27 10.9 

316 1.0 1.78 368 56 25.9 

318 4.0 3.22 370 25 14.5 

320 7.0 3.84 372 29 6.55 

322 6.0 4.25 374 12 4.34 

324 3.0 4.87 376 17 3.07 

326 5.0 2.98 378 12 3.65 

328 11.0 6.72 380 22 5.28 

330 26.0 6.84 382 21 8.66 

332 21.0 6.60 384 33 13.9 

334 14.0 6.14 386 9 9.36 

336 12.0 4.87 388 5 4.13 

338 27.0 6.22 390 2 2.02 

340 24.0 7.42 392 1 1.44 

342 26.0 18.6 394 < 0.4 0.86 

344 18.0 6.82 * 396 <0.4 0.67 

346 20.0 7.32 398 <0.4 0.41 

348 23.0 5.54 400 < 0.4 0.29 

(a) Cox and Derwent (1976). 

(b) Graham (1975) (original vaIues X 2A). 
36 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Table 10 

C10 3 Absorption Cross Sections from 200 to 350 Nanometers 

x 1018 a x 08o 

(can 2 ) (nmr) (rim) (cm 2 ) 

200 5.3 280 4.6 

210 5.0 290 4.3 

220 4.8 300 4.0 

230 4.3 310 3.2 

240 3.5 320 2.5 

250 3.7 330 1.8 

260 4.3 340 1.1 

270 4.5 350 0.76 

Table 11 

HOCi Absorption Cross Sections 

x 1020 a x 10200 

(imn) (cm2) (nm) (cm2) 

290 6.3 380 2.4 

300 9.5 390 1.6 

310 11.5 400 0.8 

320 13.5 410 0.6 

330 11.9 420 0.5 

340 9.5 430 0.4 

350 7.9 440 0.3 

360 4.8 450 0.2 

370 2.8 

yield for decomposition has not been measured, the continuous nature of 

the spectrum indicates that it is likely to be unity. 

HOCI + hp -> OH + Cl (Reaction 27) 

The ultraviolet absorption cross sections of HOCI have been measured by 

Ferguson et al. (1936) and, more recently, by DeMore (1976). The recommended values, listed in table 11, are taken from the data of DeMore 
LABORATORY MEASUREMENTS 37 

(1976). The photodecomposition products have not yet been experimentally determined, although Cl atoms and OH radicals appear most 

likely. A theoretical study by Jaffe (1976) indicates that HCI formation 

is quite improbable. 

ClNO-+hv - Cl+ NO (Reaction 28) 

Nitrosyl chloride-a green gas-has a continuous absorption extending 

beyond 650 nanometers. There is good agreement between the work of 

Martin and Gareis (1956) for the 240- to 420-nanometer wavelength 

region, of Ballash and Armstrong (1974) for the 185- to 540-nanometer 

region, and of ilies and Takacs (1976) for the 190- to 400-nanometer 

region. These results indicate that the early data of Goodeve and Katz 

(1939) were seriously in error between 186 and 300 nanometers, whereas, 

at longer wavelengths, they are in good agreement with the more recent 

measurements. 

The results of Ballash and Armstrong (1974) and of lilies and Takacs 

(1976) are listed in table 12. The two sets of measurements agree within 

20 percent, except in the region near 240 nanometers, where the values of 

Ballash and Armstrong are about 60 percent higher. The recommended 

cross sections (also listed in table 12) were obtained by taking the mean of 

the two studies. 

The quantum yield for the primary photolytic process (reaction 28) has 

been reviewed by Calvert and Pitts (1967); it is unity over the entire visible 

and near-ultraviolet bands. 

C/NO2 + hp ->Products (Reaction 29) 

The absorption cross sections of nitryl chloride, C1NO 2 , have been measured between 230 and 330 nanometers by Martin and Gareis (1956) and 

between 185 and 400 nanometers by lilies and Takacs (1976). The results 

are in good agreement. Table 13 lists the recommended cross sections, 

taken from liies and Takacs (1976). 

The photochemistry of CINO 2 has not yet been studied. Likely photolysis 

products are Cl and NO2 , and the quantum yield for decomposition is 

probably unity, due to the characteristics of the spectrum. 
38 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Table 12 

CINO Absorption Cross Sections from 190 to 400 Nanometers 

x a (cm 2 ) 

(nm) (a) (b) Mean 

190 5.34(-17) 5.20(-17) 5.27(-17) 

200 7.19 6.74 6.97 

210 3.36 2.99 3.18 

220 1.26 1.07 1.17 

230 4.36(-18) 3.17(-18) 3.77(-18) 

240 1.91 7.68 (-19) 1.34 

260 1.99 (-19) 1.61 1.80 (-19) 

280 1.13 9.35 (-20) 1.03 (-20) 

300 1.03 8.67 9.49 

320 1.33 1.08(-19) 1.21 (-19) 

340 1.50 1.24 1.37 

360 1.30 1.13 1.22 

380 8.86 (-20) 7.78 (-20) 8.32 (-20) 

400 5.12 5.15 5.14 

(a) BaUash and Armsrong (1974). 

(b) Uilies and Takacs (1976). 

Table 13 

CINO2 Absorption Cross Sections from 190 to 400 Nanometers 

(nn) (cm2) (nm) (cm2) 

190 2.69 (-17) 300 1.54 

200 4.55(-18) 310 1.32 

210 3.39 320 1.02 

220 3.42 330 7.11 (-20) 

230 2.36 340 4.81 

240 1.40 350 3.06 

250 9.85 (-19) 360 J-82 

260 6.37 370 1.07 

270 3.73 380 0.62 

280 2.31 390 0.38 

290 1.80 400 0.21 
LABORATORY MEASUREMENTS "39 

CIONO 4-hp -- Products (Reaction 30) 

Measurements in the near-ultraviolet of the cross sections of chlorine 

nitrite (C1ONO) have been made by Molina and Molina (1977). Their 

results are listed in table 14. The characteristics of the spectrum and the 

instability of ClONO strongly suggest that the quantum yield for decomposition is unity. The CL-O bond strength is only about 20 kilocalories, so 

that chlorine atoms are likely photolysis products. 

Table 14 

CIONO Absorption Cross Sections at 231 K 

1020 x o X 1020 U 

(aIm) (cm2) (nri) (cm2) 

235 215.0 320 80.3 

240 176.0 325 75.4 

245 137.0 330 58.7 

250 106.0 335 57.7 

255 65.0 340 43.7 

260 64.6 345 35.7 

265 69.3 350 26.9 

270 90.3 355 22.9 

275 110,0 360 16.1 

280 132.0 365 11.3 

285 144.0 370 9.0 

290 144.0 375 6.9 

295 142.0 380 4.1 

300 129.0 385 3,3 

305 114.0 390 2.2 

310 105.0 395 1.5 

315 98.1 400 0.6 

CIONO 2 +hp -+ Products (Reaction31) 

The ultraviolet spectrum of chlorine nitrate has been examined in the 215­

to 300-nanometer wavelength region by Martin and Gareis (1956) and in 

the 185- to 460-nanometer region by Rowland et al. (1976) and Birks et al. 

(1976a). The last two sets of Aieasurements are in excellent agreement 

with each other, as well as with the earlier measurements in the shorter 

wavelength region. Table 15 lists the data of Rowland et al. (1976). 
40 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Table 15
 

CIONO2 Absorption Cross Sections from 186 to 460 Nanometers
 

X 1020 a X 1020 U 

(nm) (cm2) (nm) (cm2) 

186 995.0 300 3.91 

190 690.0 305 2.79 

195 502.0 310 2.03 

200 372.0 315 1.45 

205 344.0 320 1.07 

210 348.0 325 0.79 

215 375.0 330 0.61 

220 376.0 335 0.48 

225 307.0 340 0.38 

230 231.0 345 0.34 

235 159.0 350 0.29 

240 118.0 360 0.23 

245 85.4 370 0.19 

250 65.7 380 0.15 

255 50.9 390 0.11 

260 40.7 400 0.085 

265 32.8 410 0.059 

270 26.1 420 0.042 

275 20.2 430 0.028 

280 14.5 440 0.019 

285 10.5 450 0.013 

290 7.34 460 0.008 

295 5.12 

Although the primary quantum yield for photodissocialion of CIONO2 has 

not yet been measured, it is very likely to be unity because of the continuous nature ofthe spectrum and the large excess of available energy over 

that needed for decomposition. Possible photolytic channels include the 

formation of CIO + NO2 , Cl + NO3, and ClONO + 0. The quantum yield 

for disappearance of CIONO 2 around 300 nanometers is about 4 for 

samples of pure CIONO 2 at pressures between 10 and 100 torr (Smith et 

al., 1977), a result that can be best explained by assuming that the primary 

quantum yield for photodecomposition is indeed unity. 
LABORATORY MEASUREMENTS 41 

Chlorinated Hydrocarbons 

The primary process in the photodissociation of chlorinated hydrocarbons 

is well-established: absorption of ultraviolet radiation in the lowest frequency band is interpreted as an n-a* transition involving excitation to a 

repulsive electronic state (antibonding in C-C), which dissociates by breaking the carbon-chlorine bond (Majer and Simons, 1964)_ As expected, the 

chlorofluoromethanes, which are just a particular type of chlorinated 

hydrocarbons, behave in this fashion (Sandorfy, 1976). At shorter wavelengths, two halogen atoms can be released simultaneously in the primary 

process. The quantum yield for photodissociation of chlorinated hydrocarbons is therefore expected to be unity; the studies that specifically show 

that this is the case for CF2 C12 , CFCI3 , and CCI4 have been reviewed by 

Watson (1977). Only some recent measurements that pertain to the 

absorption cross sections of some chlorofluorocarbons and their temperature dependence are discussed here. 

CFC13 # hv -) Products (Reaction 35) and CF2 CI2 + h -+ Products 

(Reaction 36) 

The ultraviolet spectra of the chlorofluoromethanes have been examined 

by several groups. The results are in excellent agreement, as shown in 

table 16 for CFC13 (trichlorofluoromethane, or fluorocarbon 11) and 

table 18 for CF2 CI12 (dichlorodifluoromethane, or fluorocarbon 12), which 

include the room-temperature data of Chou et al. (1976a),.Robbins et al. 

(1975), and Bass (private communication, 1976). The preferred value is 

the mean,listed in the last column of each table. The low-temperature data 

of Chou et al. (1976a) and Bass (privatecommunication, 1976) are shown 

in table 17 for CFC 3 and table 19 for CF 2CI 2 ; the agreement is also very 

good. For simplicity, the recommended low-temperature values for 

CF.C12 are given by the following expression (Chou et al., 1976a): 

-T = 0296 exp [3.6 X 104 (N - 184.9) (T - 296)] 

where 0296 are cross sections, cm 2 molecule, at 296 K (i.e., the last 

column in table 18);'k is the wavelength in nanometers; and T is the temperature in Kelvin. 

Although no such simple expression is available for CFC13 , the temperature effect at stratospherically important wavelengths (near the 200­

nanometer "window") is much smaller than for CF2C12. 
42 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Table 16
 

CC13 F Absorption Cross Sections at 298 K
 

1020 U(cm 2 ) 

(am) (i0 3 Cau1) (a) (b) (c) Mean 

186.0 54.0 255.0 247.0 226 243.0 

187.8 53.5 227.0 218.0 206 217.0 

189.6 530 197.0 186.0 175 186.0 

1914 52.5 164.0 160.0 152 1590 

193.2 52.0 141.0 134.0 125 133.0 

195.1 5115 115.0 112.0 105 111.0 

197.0 51-0 93.2 93.3 84.5 90.3 

199.0 505 74.3 74.2 70.6 73.0 

201.0 50.0 590 58.0 55.0 57.3 

203.0 49.5 45.7 44.0 46.0 45.2 

205.1 49.0 34.1 32.9 33.0 33.3 

2073 48 24.8 23,0 23 8 23.9 

209.4 48.0 17.3 16.2 16.8 16.8 

211 6 47.5 11.6 11.2 11.8 11.5 

213.9 470 7.8 7.3 7.8 7,6 

216.2 46,5 4-9 5.0 5.0 

218.6 46.0 3.0 3.2 3.1 

221.0 45.5 2.1 1.9 2.0 

223.5 450 1.2 1.2 1.2 

226.0 44.5 0.8 0.8 0.8 

44.0 

(a) Chouet al (19 7

6 a) 

(b) Robbins et al. (1975). 

(c) Bass (private communication, 1976). 

CF3CI + hp -+ Products (Reaction 37) and CC12F CCl 2 + hv Products 

(Reaction 42) 

The absorption cross sections of CF3 C1 (fluorocarbon 13) and CC12 FCCF, 

(fluorocarbon 113) have been measured at room temperature by Chou et 

al. (1976b); the results are listed in table 20. 

CCIF2C CIF2 + hy -) Products (Reaction 43) and CCIF2 CF + hp -+Products 

(Reaction 44) 

Two groups (Chou et al., 1976b; and Robbins, private communication, 

1976) have examined the spectra of CCIF 2 CCIF 2 (fluorocarbon 114) 

and CClF2 CF3 (fluorocarbon 115). Table 21 lists the results; the recommended value is the mean, which is also listed in table 21. 
LABORATORY MEASUREMENTS 43 

Table 17 

Low-Temperature Absorption Cross Sections for CCI3F 

x 1020 a(cm 2) 

(urm) (I03 em "1) 252 K (a) 232 K (a) 213 K(a) 222 K (b) 

186.0 54.0

53.5 2330 

187.8 530 202.0 

189.6 525 1760 

191.4 520 164.0 161.0 161.0 143,0 

193.2 51 5 141.0 137.0 137 0 1200 

195.1 51.0 114.0 110.0 110.0 970 

1970 505 91.3 88.5 885 798 

199.0 50.0 721 69.1 691 626 

201.0 49.5 56,6 543 53.1 50 8 

203.0 49.0 43,0 41.1 40.2 38.9 

205.1 485 31.7 30.0 286 28.6 

207.3 48.0 22.6 21.1 19.8 19.6 

209.4 475 15.2 14.2 13.3 135 

211,6 470 9.9 9.1 8.5 87 

213,9 46.5 6A 5.7 54 

216.2 46.0 3 9 3.4 33 

218.6 45.5 2.3 2.0 19 

221.0 450 12 

2235 44,5 0.68 

226.0 440 0.40 

(a) Chou et al (1976a). 

(b) Bass (private communication, 1976). 

CH3 CCI 3 + hv -+ Products (Reaction 45) 

The absorption cross sections of trichloroethane, CH 3 CC13 , are listed in 

table 22. The data are taken from Rowland (private communication, 

1976). Christiansen et al. (1972) have studied the photochemical decomposition of this molecule in air. By analysis of the reaction products, 

the quantum yield for phosgene formation was determined to be 1.3, and 

the quantum yield for the primary process was assumed to be unity. An 

absorption spectrum from 190 to 220 nanometers was also reported in 

graphical form, but it is not suited for quantitative purposes. No information pertinent to this spectrum was provided, and the results are in poor 

agreement with those listed in table 22. 
44 CILOROFLUOROMETHANE WORKSHOP REPORT 

Table 18 

CC12 F2 Absorption Cross Sections at 298 K 

10,

0 o(crn2 ) v 

- (nn) (I 0o cm ) (a) (b) (c) Mean 

186.0 54.0 105.0 108.0 1040 106.0 

187.8 53,5 86.5 84.5 85,1 85.4 

1896 53.0 66.1 62.8 64.8 64.6 

191.4 52.5 51A 46.0 48.7 487 

193.2 52 0 36.7 34.0 35.3 35.3 

195 1 515 24.9 24.3 24.3 24.5 

197.0 51.0 168 16.5 166 16.6 

199.0 50.5 10.9 11.0 10.5 10.8 

201.0 &96 7.00 6.65 6 87 

203.0 49.5 4.37 4.40 4.32 4.36 

205 1 49.0 2.66 2,60 2.52 2 59 

2073 48.5 1.53 1.50 1,47 1.50 

209.4 48.0 0.90 0.87 0.90 0.89 

2116 47.5 0.53 0.49 0.50 0.51 

213.9 47.0 0,32 0.27 0.28 0.29 

2162 46 5 0.19 0.17 0.16 0.17 

218,6 0 0.10 0.089 0.095 

221 .0 0.05 0.05 

223.5 45.0 < 0.05 < 0.05 

2260 44.5 < 0.05 < 0.05 

44.0 

(a) Chou etal.(1976a) 

(b) Robbins et aL (1975).

(c) Bass (private communication, 1976). 

Chlorofluorocarbonyls 

CCI20 + hv -) Products (Reaction 46), CCIFO +1h -* Products 

(Reaction 47), and CF2 0 + hv ->Products (Reaction 48) 

Table 23 shows the absorption cross sections of CC2 0 (phosgene), CFC1O, 

and CF20 taken from the work of Chou et al. (1977). The spectrum of 

CF2 0 shows considerable structure. The values listed in table 23 are 

averages over each 50-wavenumber interval. Preliminary photochemical 

studies (Chou et al., 1977) indicate unit quantum yield for photodissociation at 184 nanometers. The spectrum of CFCIO shows less structure, and 

the CC120 speetrumis a continuum; its photodissociation quantum yield is 

unity (Calvert and Pitts, 1967). 
LABORATORY MEASUREMENTS 45 

Table 19
 

Low-Temperature Absorption Cross Sections of CC 2F2
 

1020 U(Cm2) 

(nm) (103 tmi1) 252 K Ca) 230 K (a) 222 K (b) 

186.0 54.0 103.0 102.0 99.7 

187.8 

189.6 191.4 

53.5 

53.0 52.5 

83.0 

61.5

46.3 

80.4 

58.8 44.2 

81.7 

60.7 42.1 

193.2 52.0 32.3 30.1 29.3 

195.1 51.5 21.2 19.4 19.2 

197.0 51.0 13.9 12.6 12.4 

199.0 50.5 8.72 7.85 7.49 

201.0 50.0 5.36 4.80 4.39 

203.0 49.5 3.28 2.84 2.58 

205.1 

207.3 

209.4 

49.0 

4&5 480 

1.94 

1.07 0.61 

1.65 

0.90 0.50 

1.46 

0.76 0.44 

211.6 47.5 0.35 0.28 0.25 

213.9 216.2 216.2 

47.0 46.5 46.0 

0.21 

0.12 

0.16 

0.09 0.05 

0.13 

0.07 0.04 

218.6 45.5 0.06 

(a) Chou et al. (1976a). 

(b) Bass (private communication, 1976). 

HETEROGENEOUS CHEMISTRY 

The existence, of aerosol particles in the stratosphere raises the possibility 

that heterogeneous processes may perturb the gas-phase chemistry. Much 

of the existing information concerning the properties of stratospheric 

aerosols has been summarized in CIAP Monographs I and IIIThe situation with regard to chemical effects is characterized in general 

by a lack of specific rate data for possible surface reactions, owing largely 

to the difficulty of carrying outsuch experiments and the fact that surface 
46 CHLOROFLUOROMETHANE WORXSHOP REPORT 

Table 20
 

Absorption Cross Sections for CCIF3 and CC12FCCIF2
 

1020 a (cm) 

x V 

(n-) (103 cm ­1 ) CCIF3 CCI 2FCCIF 2 

184.6 54.0 0.36 116.0 

186.0 0.31 105.0 

187.8 189.6 53.0 0.23 0.168 85.0 68.9 

191.4191.4~ 52.5 52.00.9 0.126 53.8 

193.2 

195.1 197.1 

51.5 

5100.064 51.0 

0.090 

0.041 

41.0 

30.0 21.3 

199.0 50.5 0.026 14.9 

199.0 ~ 201.0 

203.0 

50.0 

49.5 49.0 

0.017 

0.012 

1. 

10.4 

7.0 

205.1 48.5 4.7 

207.3 48.0 48.02.05 

209.4 47.505 

211.6 47.01.26 

213.9 

216.2 

46.5078 40.47 46.0 0.47 

218.6 4550.29 

221.0 45:00.18 

223.5 44.5 0.1t 

properties of the aerosols are not well-defined. Fortunately, however, 

most of the hypothetical surface effects can be ruled out by upper limit 

arguments. Three categories of possible surface effects are reviewed in 

the following subsectionsREMOVAL OF CHLOROFLUOROMETHANES (CFM'S) 

This question was considered in detail in the NAS report. To summarize, 

it was shown that no known heterogeneous processes, including absorption, 

adsorption, or chemical reaction, can account for significant rates of CFM 

loss due to aerosols. 
LABORATORY MEASUREMENTS 47 

Table 21
 

Absorption Cross Sections for CC1F2CC1F2 and CC1F2CF
 

102, a (cm 2 ) . 

CCF 2 CCIF 2 	 CCIF 2 CF3 

3 "1 (Pim) (10 cyn ) (a) (b) Mean (a) (b) Mean 

186,0 54.0 100 10.5 100 067 0.54 061 

187.8 54 7.71 8.10 791 0.58 0.40 0.49 

189.6 53.0 5.84 610 5.97 044 0.28 0.36 

1914 52.5 4.36 4.52 444 0.33 0.20 027 

193.2 3.18 3.08 3.13 0.24 0 15 0.20 

195,1 51.5 281 222 252 0.17 0.10 0.14 

197.0 	 510 1.44 1,63 1,54 0.11 0.075 0.093 

199.0 50.5 097 1.03 1.00 0077 0.055 0066 

201 0 50.0 066 0.67 0.67 0.050 

203.0 	 49,5 043 0.45 044 0.032 

205.1 	 49.0 0.28 031 030 

207.3 	 48.5 0.18 018 0.18 0.012 

209.4 	 480 012 0,10 0.11 

211.6 	 475 0070 0.055 0.063 

213.9 	 470 0.044 0.041 0.043 

216.2 	 46.5 0.027 

2186 	 460 0015
 

455
 

(a) Chou et al.(1976b). 

(b) Robbins (private communication, 1976). 

REMOVAL OF CIX 

In principle, some of the more reactive forms of chlorine, such as HCl, 

CIO, or CI, might be removed at significant rates from the stratosphere by 

incorporation into aerosols. There is evidence, however, that this is not 

the case. Wofsy (private communication, 1976) has shown that, on the 

basis of vapor-pressure data, HC1-H2 0 aerosols are not stable under stratospheric conditions. Similar results probably hold for HCI sulfuric-acid 

aerosols. Therefore, it is not expected that significant fractions of stratospheric HCI would be present in aerosol droplets. This conclusion is 

supported by measurements of aerosol composition that indicate low or 

zero chloride content (Lazrus et al., 1975; and Bigg et al., 1974). 

For a typical aerosol concentration of 1 per cubic centimeter (assuming 

03-micrometer radius and 2-gram-per-cubic-centimeter density), the aerosol mass concentration is about 2 X 10-13 grams per cubic centimeter. 
48 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Table 22 

Absorption Cross Sections for CH3CC13 

P 1020 U X P 100a 

(nm) (10 3 erf') (cm 2) (nm) (103 cm I

') (CM2) 

186.0 54.0 325.0 205.1 49.0 485 80.5 

187.0 284-0 207.3 480 63:9 

189.6 53.0 246.0 209.4 47.5 51.1" 

191.4 193.2 52.5 215.0 189.0 211.6 213.9 47.0 46.5281 39.4 

195.1 51.5 168.0 216.2 46.0 19.6 

197.0 51.0 148.0 218.6 45 12.5 

199.0 50.5 128.0 221.0 45.0 8,3 

201.0 500 111.0 223.5 5.1 

203.0 49.5 49 0 95.4 226.0 44.0 4. 2.9 

This quantity is comparable to the mass concentration of Cl. in lower 

regions of the stratosphere. On the basis of the aerosol composition 

measurements mentioned above, it follows that significant fractions of the 

stratospheric Clx are not present in the aerosol particles. 

Of course, even a very small aerosol chloride content could correspond to 

large loss rates if there were a rapid flux of aerosols out of the stratosphere. However, the residence time of aerosols is comparable to that of 

gaseous species. It is therefore unlikely that aerosols provide asignificant 

sink of Cl in the stratosphere. x 

CATALYTIC PROCESSES 

Aerosols might also perturb stratospheric chemistry by surface catalysis of 

reactions-that are slow in the gas phase. For an aerosol concentration of 

1 per cubic centimeter, the collision frequency for gaseous species with 

the surface is of the 6rder of 10's se -1 , with some uncertainty due to 

the size and geometry of the particle. (In most regions of the stratosphere, 

the aerosol density is probably less than I per cubic centimeter, so that - the figure 10 5 sec "1 is an upper limit.) The frequency of 10-5 see 1 thus 

constitutes a reference value against which the rates of gas-phase processes 

are to be compared. If, as is often the case, the gas reaction that accomplishes the same net effect is much faster,the heterogeneous process can be 

neglected. A possible example of a significant surface process would be 

the catalytic conversion of active forms of chlorine (C1 or CIO) to inactive 
LABORATORY 	MEASUREMENTS 49 

Table 23 

Absorption Cross Sections for CC12 0 , CCIFO, and CF2 O 

1020 a (CM2) 

(nm) (103 cm"') CC2 0 CCIFO CF 2 0 

184.9 

186.0 

187.8 

189.6 

191.4 

193.2 

195.1 

197.0199.050.5 

199.0 

201.0 

203.0 

205.1 

207.3 

209.4 

211.6 

213.9 

216.2 

218.6 

221.0 

223.5 

226.0 

54.0 204.0 4.7 

189.0 15.6 5.5 

137.0 14.0 5.2 

53.0 117.0 13.4 4.5 

52.0 93.7 12.9 4.0 

52.0 69.7 12.7 3.3 

51.0 52.5 12.5 2.8 

50.5 41.0 12.4 2.3 

31.8 12.3 1.9 

49.5 25.0 12.0 1.4 

49.0 20.4 11.7 1.1 

16.9 11.2 0.86 

15.1 10.5 0,65 

47.0 47.5 13.4 9.7 0A8 

47.0 12.2 9.0 0.36 

46.5 11.7 7-9 0.26 

46.0 11.6 6.9 0.21 

11.9 5.8 0.15 

45.0 12.3 4.8 0.12 

45 12.8 4.0 0.10 

13.2 3.1 0.08 

reservoirs (such as HC1). It can be shown, however, that these processes are 

unlikely to be'important. The frequency of Cl conversion to HCI by the 

gas-phase reaction, Cl + CH 4 -# ELCI + CH 3 , is about 10-2 see- under 

stratospheric conditions. Thus, the surface process could not possibly 

compete with the homogeneous reaction, even in the unlikely case of unit 

collision efficiency. 

For CIO, the situation is slightly different because the frequency for ClO 

conversion to 	HCI (by the sequence of reactions CIO -+ Cl >.- HC) is 

t about 10-6 see -

' This figure is obtained by multiplying the frequency of 
50 CHLOROFLUOROMETHANE WORKSHOP REPORT 

CIO conversion to Cl (about 10-2 sece through reaction with 0 and NO) 

by the fraction of those Cl atoms that react with CH4 rather than with 

03 (about 1 in 104). Since this rate is comparable to the surface collision 

rate, an appreciable effect might arise in the event of a high collisional efficiency for reduction of CI0 to HC1- However, as pointed out in the NAS 

report, CIO is known to survive 104 to i0 5 collisions with aqueous acidic 

surfaces similar to those expected for the sulfate aerosols. Therefore, there 

is no reasonable expectation that C10 would be converted to HCl on 

stratospheric aerosols at a rate sufficient to perturb the gas-phase 

chemistry. 

Perhaps heterogeneous conversion of C10 to HOC1 should also be considered, although there is no evidence that the efficiency of such a process 

would be high. Moreover, HOCt is not an effective reservoir for stratospheric chlorine because it photolyzes with a J-value of the order of 103 

see -IA surface process that may play a significant role in the stratosphere is the 

hydrolysis of chlorine nitrate. This compound can be considered to be the 

mixed anhydride of nitric and hypochlorous acid. The reaction, CINO3 + 

H2 0 -+HOC1 + HNO3 , might have a high collision efficiency on the surface of aerosol particlesin which case the collision frequency of 10- sec' 

would be competitive with the photolysis of chlorine nitrate, which has a 

J-value of about 5 X 10-5 sec*(Rowland et at., 1976). The formation of 

HOCi by this route would probably not be significant for reasons already 

given. However, production of HNO 3 by this path could be competitive 

with the gas-phase process, OH + NO2 + M -) HNO 3 + M, as can be shown 

by approximate calculations. For these reasons, additional experimental 

studies of the rate of surface hydrolysis of chlorine nitrate (and N2 05) 

are needed. 

The stratospheric H20 2 concentration may be reduced by surface destruction if the collision efficiency is high. This follows from the fact that the 

frequency of H202 destruction by reaction with OH (the dominant gas 

path) is about 106 see", which is lower than the maximum surface collision frequency. 

Loss of odd oxygen by surface destruction of either atomic oxygen or 

ozone cannot compete with the gas-phase rates at which these species are 

converted to 0,2 . Sirlarly, surface loss of OH or HO2 cannot compete 

with homogeneous termination reactions of these species. 
N77=316961
 

CHAPTER 2 

MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 

INTRODUCTION 

Variations in atmospheric ozone result from a number of interacting 

photochemical and meteorological processes; therefore, changes in the 

concentrations of some of the gases that are active in the ozone production/destruction cycle could significantly alter the amounts of ozone at 

various levels in the atmosphere and thus produce a net change in the 

total-ozone amount in a vertical column. It has been suggested that 

chlorofluoromethanes (CFM's) that are transported upwards from the 

troposphere to the stratosphere act as a significant agent for such a destruction process. It is shown in Chapter 4 of this document that the 

largest reduction of the ozone concentration (about 60 percent) resulting 

from CFM input to the stratosphere would be at about 40 kilometers 

with a net decrease of aubout 3 percent in the total amount in a time 

frame of about 40 to 50 years. It is therefore of particular interest to 

determine the natural ozone variations that occur above the ozone maximum where the most pronounced effects of the CFM's are expected to 

be manifested, as well as to determine natural changes in total ozone that 

have been observed to occur during time intervals of several years or more. 

The total amount of ozone in a unit vertical column (usually referred to 

as "total ozone") is of the order of 300 milliatmosphere centimeters 

(m-atm-cm) or 0.3 centimeter at standard temperature and pressure. This 

amount varies typically from about 250 m-atm-cm near the Equator to 

about 400 m-atm-cm at high latitudes. There are large daily and 

seasonal variations, particularly at mid-to-high latitudes where day-to-day 

changes can be as high as 40 m-atm-cm per day, and the seasonal variation 

can be about 140 m-atm-cm- The mean surface-ozone concentration in 

clear air is about 3 parts per hundred million by volume (pphrnv). Surfaceozone variations do not indicate a significant diurnal cycle, although the 

annual cycle is prominent at all latitudes. Observations of the vertical 

ozone distribution show maximum ozone concentrations at about 25 

kilometers over the Equator and at 16 to 18 kilometers in polar regions 

during the spring. There is a strong correlation between the total-ozone 

amount and the level and strength of the maximum concentration. Large 

total amounts are associated with large values of the maximum concentration situated at low atmospheric levels. 

51
 
52 9,_.tCHLOROFLUQROMETHANE WORKSHOP REPORT 

Pronounced seasonal variations are present in the lower stratosphere and 

in the middle and upper stratosphere above the ozone maximum, particularly at high latitudes. The existence of a quasi-biennial oscillation in 

ozone has also been well documented, and it has been suggested that 

there may be a solar-cycle variation. Analysis of long-term ozone data at 

a number of observing stations does not indicate any clearly defined 

relationship between volcanic eruptions andlor nuclear explosions and 

ozone variations. Discussions of the distribution of ozone and its "natural" variations are contained in the Climatic Impact Assessment Program 

(CAP, Volume 1, 1975). 

Knowledge of the data base, including the various techniques of observations in current and planned use, is essential for detecting anthropogenically caused ozone variations. This chapter briefly discusses the techniques that have been used for total-ozone and vertical distribution, 

including measurement accuracy arid precision. A description of data 

availability is provided, and sample total-ozone and stratospheric data are 

presented that illustrate the trends that have been deduced from the data. 

Comments on the problem of the analysis of ozone variations are contained in the final subsection. 

MEASUREMENTS 

TOTAL OZONE 

Ground-Based Measurements 

Observations of total ozone are currently made at about 75 locations 

over the Earth, with the majority (85 percent) of the stations located in 

the Northern Hemisphere. Dobson ozone spectrophotometers (Dobson, 

1957) are used at 53 of the stations, while other instruments, generally 

broad-band filter ozonemeters (Gushchin, 1963) are used at 22 stations, 

principally in the Union of Soviet Socialist Republics (USSR). Other 

types of instruments are used at six stations. Two new types of instruments are presently being developed for total-ozone measurements: a 

grating ozone spectrophotometer (Brewer, 1973) and a narrowband filter 

ozonemeter (Matthews et al., 1974). 

The Dobson spectrophotometer is a prism double monochromator that 

yields atmospheric total-ozone data from measurements of relative intensities of pairs of direct Sun or zenith sky scattered solar radiation at wavelengths in the 300- to 340-nanometer wavelength region where ozone 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 53 

absorption changes greatly with wavelength. In principle, measurements 

on single wavelength pairs, designated as the A, B, C, and D pairs,, are 

possible. However, in practice, observations are generally made on 

double-pair wavelengths, such as AD or CD, because, in this way, the 

effect on the measurements of light scattering by particulates is minimized. Of the various instrument types, Dobson spectrophotometers are 

currently considered to yield the most reliable data. Performance of the 

USSR filter ozonemeters has been discussed elsewhere (Bojkov, 1969; 

and Gushchin, 1972). Gushehin (1970) has described. improvements 

made to the meters since 1969, and a program is currently underway in 

Boulder, Colorado, to assess the performance of a USSR filter ozonemeter in relation to measurements made with a United States' (U.S.) 

Dobson speetrophotometer. 

The uncertainty (± lu) associated with the absolute calibration (accuracy) of Dobson spectrophotometers for AD wavelength observations on the 

direct Sun is estimated to be ±5 percent and stems primarily from uncertainties in the ozone absorption coefficients used in reducing the 

observational data. Systematic errors may be incurred through assumptions made about aerosol scattering effects and the constancy of the 

extraterrestrial solar flux (L0 ) at the spectrophotometer wavelengths. 

Because observations on different wavelength pairs have been found to 

yield somewhat different results, the International Ozone Commission 

has adopted AD direct-Sun observations as standard, since absorption 

coefficients for AD wavelengths are believed to be the least in error. 

For Dobson measurements made on clear or cloudy zenith skylight, relative uncertainties associated with the data compared to AD direct-Sun 

measurement data uncertainties are estimated to amount to < 3 percent 

for 90 percent of the observations (Komahyr, 1961) if adequate care has 

been taken in devising empirical charts for reduction of the zenith-sky 

data. Such charts are ordinarily constructed from a substantial body of 

quasi-simultaneously obtained direct-Sun and zenith-sky data. Because 

the spectral composition of light scattered from the zenith sky depends 

on atmospheric aerosol loading, ozone vertical distribution, and ground 

albedo (which may all vary at a particular station during the course of a 

year), zenith-sky observations do not yield as reliable total-ozone data as 

do direct-Sun measurements. 

Within the operational Dobson station network, larger discrepancies than 

those previously indicated have been found in comparisons among different instruments. For example, during an intercomparison of Eastern 

European Dobson instruments held in Hungary (Gushehin, 1972), differences in calibration of up to 20 percent were observed. An inter­
54 CHLOROFLUOROMETHANE WORKSHOP REPORT 

national comparison of Dobson spectrophotometers held in Poland 

(Dziewulska-Losiowa and Walshaw, 1975) revealed calibration differences 

among 10 instruments of up to 17 percent, although, for 7 of the 10 

instruments, the calibration levels did not differ by more than 8 percent. 

For a total-ozone-trend determination, the long-term Dobson spectrephotometer measurement precision is of relevance rather than the absolute measurement accuracy. Factors affecting the measurement precision 

are uncertainties associated with determinations of the instrument extraterrestrial constants, the transferring of calibrations from calibrated to 

noncalibrated spectrophotometers, and instrument drift corrections, as 

well as absorption and scattering coefficient errors and errors associated 

with the temperature dependence of the ozone-absorption coefficients. 

It is estimated from a root-mean-square error analysis that, for a carefully 

operated worldwide Dobson instrument network with good quality control, the maximum possible long-term measurement precision corresponds 

to uncertainties in direct-Sun measurement data of about ±1.5 percent 

and uncertainties in zenith-sky measurement data of approximately ±3.5 

percent. However, it is not likely that this can be achieved, and a practical measurement precision goal for such a station network might be 

uncertainties of ±2 to ±3 percent for direct-Sun data and ±5 percent for 

zenith-sky data. These measurement precision uncertainties would be 

essentially random if appropriate instrument drift corrections were appied when needed. Therefore, they would not contribute significantly 

to uncertainties in trend determinations from data sets consisting of more 

than several years observations. 

Within the U.S. 12-station Dobson spectrophotometer network operated 

during the past 15 years, uncertainties associated with the direct-Sun 

observations on AD wavelengths are estimated to be ±2 to ±3 percent, 

while those for zenith-sky observations are estimated to be ±5 to ±6 

percent. Improvement in the quality of the existing data body can be 

made by applying final corrections to the data and reprocessing them. 

The attainment of long-term, high-precision total-ozone measurements 

within the global total-ozone station network will require periodic (3- to 

5-year interval) intercalibrations of primary standard and secondary standard spectrophotometers and an ongoing program of intercalibrations of 

secondary standard and field instruments. At a recent meeting in 

Dresden, German Democratic Republic, the International Ozone Commission recommended that U.S. Dobson spectrophotometer No. 83 be 

designated by the World Meteorological Organization (WMO) as the 

Primary Reference Standard Instrument for total-ozone measurements 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 55 

and that Dobson spectrophotometers Nos. 41 and 64 (Europe), No. 77 

(Canada), No. 108 (USSR), No. 112 (India), and instruments from Japan, 

Australia, Africa, and South America be designated as Secondary Reference Standard Instruments. Plans are currently underway for an international comparison of secondary standard Dobson spectrophotometers 

with U.S. instrument No. 83 to be held at the National Oceanic and 

Atmospheric Administration (NOAA), Boulder, Colorado, in July 1977. 

The question of whether or not significant variations in L0 , the extraterrestrial solar flux at Dobson spectrophotometer wavelengths occur during the course of a solar sunspot cycle remains unanswered. Although 

such variations, if any, are not expected to affect the precision of totalozone measurements beyond the +l-percent level, an optimum monitoring program requires that L. variations be monitored by the standard 

Dobson spectrophotometer on a continuous basis and that any detected 

changes in L° values be transferred to the field instruments by the secondary standard spectrophotometers, as required. An ongoing program 

to monitor % variations with U.S. Dobson spectrophotometer No. 83 

should therefore be implemented at Mauna Loa Observatory as soon as 

possible. Mauna Loa Observatory is especially suited for such measurements because of the clarity of the atmosphere at that site and the relatively small day-to-day variations in total ozone that occur there during 

the course of a year. 

In the future, satellite observations will play an important role in global 

total-ozone measurements and the monitoring of secular changes in ozone 

on a global scale. The satellite-based measuring instrumentation will 

require calibration data from a select set of ground-based total-ozone 

measuring instruments. Conversely, satellite observations of total ozone 

may play a useful role-in identifying instruments within the worldwide 

Dobson speetrophotometer station network that require recalibration. 

Satellite Measurements 

Remote sensing from satellites permits observations over most of the 

globe with a uniformity of space and time sampling that is not obtainable with any other kind of observing network. Obtaining the coverage 

with a single instrument results in improved measurements of gradients 

and trends, and eliminates the requirement of intercomparing the results 

obtained from a large number of instruments. Satellite sounding is also 

cost-effective because, although the initial investment is large, the cost 

per datum is extremely small. 
56 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Backscattered Ultraviolet 

This technique utilizes measurements of the ratio of the backscattered 

ultraviolet radiation to the incident solar flux in the long wavelength end 

of the Hartley-Huggins bands. An additional measured radiance in an 

absorption-free spectral region adjacent to the band is used to determine 

the required lower boundary reflectivity, whereas a weak dependence on 

vertical distribution is accounted for in a statistical sense through the use 

of average profiles in precalculated lookup tables. Estimates of ozone are 

obtained for a lower boundary either of clouds or the planetary surface, 

with the decision and any interpolation based on the measured albedo in 

the spectral region outside the absorption band. 

Dave and Mateer (1967) have studied the sensitivity of ozone estimates to 

instrument noise and errors in lower boundary reflectivity and pressure. 

A 1-percent error in the measurement of the ratio of backscattered radiance to incident solar radiance is found to propagate into an error of 

approximately 1 percent in ozone-column amount. The major sources of 

error are associated with the necessity of modeling the lower boundary. 

Because of the wavelength dependence on the reflective properties of natural surfaces, clouds, and aerosols, the reflectance measured outside the 

absorption band must be extrapolated to wavelengths within the band. 

This can contribute to the overall error, since a 5-percent error in reflectance is found to propagate typically into ozone errors of 2 to 3 percent. 

Uncertainties due to lower-boundary pressure ,errors are found to be of 

the same order. Detailed estimates of sensitivity to other factors, such as 

aerosols and non-Lambertian and nonhomogeneous lower boundaries, 

have not yet been made; however, efforts are underway to address these 

problems. 

Comparisons between total-ozone estimates from the Nimbus-4 backscattered ultraviolet instrument (BUV) and ground-based Dobson measurements have been made. Mateer et al. (1971) analyzed a heterogeneous sample of 320 near-coincident Dobson and Nimbus4 BUV measurements and found the satellite values to be smaller than the Dobson 

values for low total ozone and larger than the Dobson values for high 

total ozone. The standard error of estimate was found to be approximately 6 percent. Similar comparisons are now being made for individual 

stations as a function of time, but results are not available. 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 57 

Nadir-Viewing Infrared 

The nadir-viewing infrared technique utilizes measurements of upwelling 

radiation within the 1042 cm ozone-absorption band. Ozone-column 

amounts are determined from the alteration of surface radiation after 

corrections are made for emission by the atmosphere itself. An important limitation of the method is the need for a strong temperature contrast between the underlying surface and the ozone layer-a condition 

that is often poorly met at high latitudes in winter. 

Prabhakara et al. (1970) compared Nimbus-3 infrared interferometer 

spectrometer (IRIS) ozone measurements with Dobson spectrophotometers for 23 stations and found a standard deviation of the differences 

of 0.020 centimeters or 5.8 percent of the mean Dobson total. A comparison of results from the higher spectral-resolution Nimbus-4 IRIS with 

Dobson measurements at Arosa, Switzerland, and Macquarie Island 

(Prabhakara at al., 1976) shows good agreement in day-to-day aone 

fluctuations with a random difference of about 6 percent. However, on a 

seasonal time scale there are systematic differences, as large as 20 percent 

at high northern latitudes. 

Programs to evaluate the overall accuracy and precision have not been 

established for either of the foregoing satellite techniques; therefore, 

at the present time, it is difficult to establish a quantitative error estimate 

for the determination of ozone trends from satellite measurements. However, as discussed in the preceding subsection, it is anticipated that, for 

analysis of ozone trends, satellite and Dobson data will be combined in a 

complementary fashion to provide results of better quality than could be 

achieved with either approach alone. Further work is required for determining the most effective way of combining these data. The desired outputs include not only the link between the two systems but also a characterization of the individual precisions and long-term stabilities. 

OZONE VERTICAL DISTRIBUTION 

Ground-Based Measurements 

Dobson spectrophotometers have been used to determine ozone profiles 

by a method developed by F. W. Gotz in 1929 (see, for instance, Craig, 

1965). This method involves observation of the ratio of the intensities of 
58 CHLOROFLUOROMETHANE WORKSHOP REPORT 

downward scattered zenith radiation at two wavelengths in the Huggins 

bands (one strongly absorbed, the other weakly absorbed) as the solar 

zenith angle varies from about 600 to 900. As the solar zenith angle 

increases, the shorter wavelength radiation is absorbed more and more 

compared to the longer wavelength radiation, but, at a zenith angle of 

about 850, this relation tends to be reversed (hence, the name Umkehr, 

which means "turnaround") because relatively more of the shorter wavelength radiation is scattered vertically downward through the ozone layer. 

An estimate of the vertical ozone distribution can be obtained from the 

ratio of the radiation intensities at the two wavelengths as a function of 

solar zenith angle. The mean ozone concentration is calculated for nine 

layers (each about 5 kilometers deep) up to about 45 kilometers. 

Umkehr observations are sensitive to the normal factors involved in all 

inversion evaluations; namely, measurement errors and errors in a priori 

statistics. In addition, however, changes in optical effects due to aerosols, 

temperature dependence of ozone absorption in the Huggins bands, and 

the effects of surface reflectivity all add to uncertainties in the application of the Unkehr method. Nevertheless, Umkehr observations represent the only long-term data base (approximately 20 years for some stations) for variations above the level of ozone maximum (see, for instance, 

Dutsch, 1974). 

Attempts to evaluate the accuracy of the Umkehr results using simultaneous direct ozone observations have shown that the Umkehr measurements 

give less ozone at the level of the major maximum and more at the upper 

and lower levels than observed by the ozonesonde. This is due, in part, to 

deficiencies in the first-guess statistics used in the inversion procedure, 

and probably to a lesser extent to the other error sources discussed previously. Evaluation of data from Tallahassee, Florida; Aspendale, 

Australia; and Arosa, Switzerland, indicate that, in the low levels, the 

average difference is on the order of 20 to 40 percent, 10 to 20 percent 

at the ozone: maximum, and 10 to 30 percent at the maximum altitude of 

the balloon observations (approximately 30 kilometers). It should be 

noted that accuracy of the balloonsonde decreases above about 25 kilometers. Vertical ozone-profile Umkehr observations are more sensitive 

to ozone changes at the higher levels (greater than 25 kilometers), and 

increasingly better results would be expected with altitude above that 

level. If random errors in the observations tend to cancel for a large number of observations, the average profiles become meaningful, and a large 

number of observations may be quite useful in determining time changes 

in gross features of the vertical ozone profile. 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 59 

Rocket, Balloon, and Aircraft Measurements 

Direct ozone soundings are a desirable part of any comprehensive program to test model predictions at the levels of greatest sensitivity to CFM 

effects (35 to 45 kilometers). These measurements can be used to determine long-term changes either directly, or indirectly used as ground-truth 

data for satellite remote sensors. Sounding instruments have been developed using optical, chemical, and chemiluminescent methods (table 24). 

OpticalMethods 

Optical methods that use the height varying differential absorption of 

sulight in the Hartley and Huggins bands of ozone are absolute. Results 

are generally reported as the integrated ozone concentration in layers 1 or 

2 kilometers thick. Because the product of height resolution and precision is proportional to the signal-to-noise ratio, precision is improved by 

degrading the height resolution. The optimum height region for optical 

ozone soundings is from 25 to 55 kilometers, where the optical depth 

is a rapidly varying function of height. Several wavelength bands can be 

used to maintain good precision, as well as to produce redundant data to 

test internal consistency. The methodology for routine optical soundings 

was devised by Kulcke and Paetzold (1957), who produced a two-band 

filter wheel radiometer for balloons. Optical methods for inexpensive 

rockets were later developed in the United States by Krueger and 

McBride (1968), in Australia by Sissons (1968), and in Japan by Nagata 

et al. (1971). The U.S. instrument uses four filters at wavelengths between 250 and 320 nanometers to obtain the dynamic range for the 

ozone changes in the height range from 25 to 60 kilometers. The measurement precision is better than 5 percent for 2-kilometer layers, and the 

accuracy is estimated to be between 5 and 15 percent. Closed opticalabsorption cells for surface-ozone measurements have been modified by 

Hilsenrath and Ashenfelter (1976) for balloon applications and by Falconer and Holdeman (1976) and Loewenstein et al. (1975) for aircraft 

use. The accuracy is potentially quite high, although the present altitude 

limit is about 35 kilometers. 

Chemical Methods 

Chemical methods are ideal for measurements in the troposphere and the 

lower stratosphere. Brewer and Milford (1960) developed an instrument 

based on the reaction of ozone with potassium iodide (KI) that is capable 
Technique 

Chemical 

Balloonsonde 

Chemiluminescence 


Balloonsonde 


Rocketsonde 


Reference 

Brewer and 

Milford 

(1960) 

Komhyr and 

Harris (1971) 

Regener 


(1960) 


Hilsenrath 

et al. 

(1969) 

Randhawa 


(1966) 

Table 24 

Direct Ozone Measurement Techniques 

Height Range 

Data Type (kilometers) Advantages 

KI bubbler 0 to 30 	 Good vertical 

resolution 

KI electro- 0 to 30 Good vertical 

chemical resolution 

concentration 

Dye chemi- 0 to 30 High sensi. 

luminescence tivity, good 

vertical 

resolution
 

Self-pumping 35 to 70 High sensi. 

dye cheim tivity; good 

luminescence vertical 

resolution 

Self-pumping 20 to 60 Good vertical 

dye chemi- resolution 

luminescence 

g
 

Disadvantages 

Height limited by 

solution boiling 

point; subject to 

contamination; air- 0 

pump errors at higlh 

altitudes, requires 

correction to Dobson 

total ozone; fieasures 

total ozidant 

0 

Calibration stability 

problems 

Vt 

Long ealibration 

procedure; cost high 

Calibration stability 

problems 
Table 24 (Continued) 

Heiglt Range 

Technique Reference Data Type (kilometers) Advantages Disadvantages 

Optical 

Balloonsonde Paetzold and Differential 15 to 30 Absolute Limited vertical 

Piscolar absorpition of resolution response, 

(1961) sunlight day only; diffuse 

Rocketsondc Krueger Differential 20 to 65 Absolute; skylight problems at heights less than 

z 

U 

(1965) absorption of internally 20 kilometers; inter- Z 

sunlight redundant ference by other 

absorbers could exist 

Sissons Differential 20 to 65 Absolute, 

(1968) absorption of internally 0 

sunlight redundant 0 

N 

Nagata et al Differential 20 to 65 Absolute; 

(1971) absorption of internally 

sunlight redundant 

Balloon/ailrcraft Ililsenratli Absorption cell 0 to 35 Absolute 

and Ashenfelter O 

(1976), Loewen- Z 

stein ct al 

(1975) 
62 CHLOROFLUOROMETHANE WORKSHOP REPORT 

of fast response and that has demonstrated repeatability within a few percent. A variation of the KI technique was developed by Komhyr and 

Harris (1971). This instrument shows better agreement with Dobson 

total-ozone data than does the Brewer instrument. 

Chemiluminescent Methods 

The use of dry chemiluininescent methods for balloon ozonesondes was 

developed by Regener (1960). The primary advantages are high sensitivity (needed for measuring tropospheric ozone), specificity to ozone, 

and inherent simplicity. Unfortunately, production instruments exhibit 

instabilities, and the technique is no longer used on balloons. The high 

sensitivity and simplicity have continued to be attractive for rocket purposes. Hilsenrath et al. (1969) and Randhawa (1966) have pioneered this 

development, which uses self-pumping into an evacuated chamber for air 

flow. By careful attention to calibration detail, Hilsenrath (1969) obtains close agreement in ozone profiles with other techniques from 25 to 

65 kilometers. He computes uncertainties of 20 percent between 25 and 

55 kilometers. 

Data Quality and Accuracy 

The WMO and the International Association of Meteorology and Atmospheric Physics (IAMAP) have acted to maintain the quality of ozone 

measurements by sponsoring instrument intercomparisons through the 

International Ozone Commission. 

Balloon Intercomparisons-In 1970, international tests compared data 

from the Brewer and Komhyr instruments, and Japanese, Indian, Italian, 

and East German (GDR) versions thereof (Attmannspacher and Dutsch, 

1970). This flight series demonstrated that the operational instruments 

generally determined the fine structure reproducibly up to at least 30 

millibars (-'25 kilometers) and that adjustment factors were required to 

obtain agreement with Dobson total-ozone amounts. 

The adjustment factors, which depended on the instrument type, were 

constant within a standard deviation of 5 to 9 percent. If the data are 

averaged over 4.5-kilometer layers, the instruments differ from the mean 

by less than 5 percent in the stratosphere, but errors of up to 25 percent 

are found in the troposphere. 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 63 

The balloonsonde intercomparisons have been productive in determining 

instrument precisions and accuracies and in eliminating unstable instruments. Contemporary production wet-chemical ozonesondes are internally consistent within 10 percent (better if coincident total-ozone data 

are available) and can detect fine structure from the surface to at least 

30 millibars. The accuracy at upper levels (25 to 35 kilometers) has not 

been tested. At these levels air-pumping errors and solution freezing can 

be significant problems. WMO and IAMAP are planning a new balloonsonde intercomparison series for 1977. 

Rocket Intercomparisons-Regularintercomparisons of different rocket 

ozonesondes have not been conducted; however, several ad hoc comparisons have been made between U.S. instruments so that relative merits 

can be assessed. In 1968, the Hilsenrath chemiluminescent and Krueger 

opticalsondes were compared and showed no differences greater than 

10 percent between 30 and 50 kilometers, including small-scale structure. 

In 1975, the latest comparisons of these instruments show agreement 

within 10 percent above 40 kilometers, but differences are greater than 

20 percent at lower altitudes. Additional comparisons of these and other 

rocket instruments would serve to resolve these differences. Prior to this, 

experimenters should be encouraged to conduct repeatability flights to 

determine internal consistency. 

The optical techniques are capable of better than 5 percent accuracy 

when properly applied at altitudes between 25 and 55 kilometers. With 

careful testing for calibration stability, the chemiluminescent techniques 

should be capable of similar accuracies. 

Rocket and Balloon Comparisons-Few comparisons have been held. 

Limited experience with Komhyr chemicalsondes and Krueger opticalsondes indicates agreement within 5 percent at 25 kilometers but 20­

percent differences at 30 kilometers and above. 

Satellite Measurements 

BackscatteredUltraviolet 

The backscattered ultraviolet technique has been employed in experiments flown on Nimnbus-4 (Heath et al., 1973), Orbiting Geophysical 

Observatory4 (OGO-4) (Anderson et al., 1969), and Atmosphere Explorer-E (AE-E). This method makes use of radiance measurements within an 
64 CHLOROFLIJOROMETHANE WORKSHOP REPORT 

ozone-absorption band at several different wavelengths corresponding to 

scattering layers that cover a range of heights in the stratosphere. "Topside." profiles can be retrieved from these data between approximately 

the 0.5- and 10-millibar levels. The information content of the measurements is limited by both the nature of the radiative transfer process and 

by the instrumental measurement errors. Analyses indicated that, for 

measurement errors of 1 percent, the mixing-ratio profile can be retrieved with a formal root-mean-square error of about 10 percent and a 

vertical resolution of approximately one pressure-scale height (about 7 

kilometers). Little profile information can be extracted below 10 millibars; in this region, the profile can probably be better estimated through 

statistical correlation with the total column abundance (Sellers and 

Yarger, 1969). 

Assumptions that are generally made in retrieving topside ozone profiles 

include molecular scattering and horizontal homogeneity. For instruments such as the Nimbus-4 BUV, in which the field of view is slightly 

offset between measurements at successive wavelengths, the profile is 

representative of spatially averaged conditions with the upper part of the 

profile representative of a region slightly different from that of the lower 

part. The assumption of molecular scattering is violated when stratospheric aerosols are present, and high-level dust layers may result in significant scattering contributions that can affect high-level retrievals 

(Elliott, 1971; and Cunnold et al., 1973). However, nearly simultaneous 

rocket and satellite retrieval profiles indicate little or no aerosol effect 

during 1970 and 1971. 

Long-term stability of profile determinations necessitates the periodic 

measurement of incident solar flux. Practically, this requires that a spectrally stable transfer device be inserted in the optical train to fill the field 

of view. If stability is in doubt, measurements of ozone profiles to above 

50 kilometers by another technique could be used to assess the solar flux. 

Infrared Limb Viewing 

Ozone distributions can be determined from measurements of the profiles of emission by the 1042-cm" band of ozone, as viewed at the 

Earth's limb (Gille et al., 1970). One of the advantages of this approach 

is the high vertical resolution (3 to 5 kilometers) that is possible. Other 

advantages are the reduction in ambiguity in data interpretation when 

viewing the emission against the cold background of space and nightside, 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 65 

as well as dayside, operation. Data have now been obtained from the 

Limb Radiance Inversion Radiometer (LRIR) (GUile et al., 1975), which 

was launched on Nimbus-6 in June 1975. 

2 From a noise of 0.01 W m" sfI determined for the LRIR, prelaunch 

calculations suggest a precision of 23 percent in the altitude range of 15 

to 21 kilometers, 5 percent between 30 and 36 kilometers, and 18 percent from 45 to 51 kilometers, with an average of 12 percent over the 

entire 15- to 51-kilometer interval. Modified inversion procedures should 

improve these results. 

A comparison of an LRIR retrieval and an optical rocket measurement 

at Wallops Island in summer (figure 1) gave a root-mean-square percentage difference of 14 percent between the satellite and rocket over the 

20- to 55-kilometer range (Gille et al., 1976a and 1976b). This value is 

very close to the precision claimed for the rocket flight, as well as to the 

overall value noted previously. The conclusion is that the LRIR retrieval 

and the rocket measurement cannot'be distinguished in quality. 

The LRIR data cover the globe from 800 N to 65°S for the 7 months of 

operation. The reduction is now proceeding rapidly, and the data look 

very promising. Further comparisons of retrievals with rockets and the 

Nimbus-4 and AE BUV results are planned. 

Occultation Measurements 

Hays et al. (1972) and Reigler et al- (1976 and 1977) have shown that 

ozone profiles above about 50 kilometers can be determined by observing the occultation of stars by the ozone atmosphere at wavelengths in 

the Hartley band. Differences among the various observational results 

and between observations and theoretical calculations for the ozone densities in the tropical mesosphere and lower thermosphere suggest the need 

for improvements in the theory and a continued program of ozone observation at these altitudes. Pepin plans to determine ozone profiles from 

solar occultations in the Chappuis band during the Stratospheric Aerosol 

and Gas Experiment (SAGE), planned for a 1978 launch. Studies of 

occultation determination in the near-infrared have been performed by 

Drayson et al. (1972). Because no extensive body of occultation data 

exists at this time, no further discussion of this technique will be given. 
66 CHLOROFLUOROMETHANE WORKSHOP REPORT 

70 i 

65 

60 0 KRUEGER 

I HILSENRATH 

SS 

50 +0 500 

+ 0 

E 45 

= +0 

20 00 

35 0
 

0
 

o0 

02 0 

25 0 

20
 

is +


00O 4 6 8 To 12 

OZONE VOLUME MIX[( RATIO OPPMV) 

Figure 1. Ozone profile in the stratosphere over Wallops Island, Virginia, 

on July 29, 1975, determined by three different techniques. The solid 

line is a preliminary retrieval from measurements by the LRIR aboard the 

Nimbus-6 satellite. Symbols show in situ measurements from rockets by 

the Krueger optical ozonesonde and the Hilsenrath chemiluminescentsonde. 

Additional Measurement Techniques 

Several additional measurement techniques either have been proposed or 

are under development. Because no body of data applicable to the CFM 
MEASUREMENTS AND ANALYSIS-OF OZONE VARIATIONS 67 

problem has been produced by these techniques, they will not be discussed in detail. For completeness, differential absorption of laser radiation, laser heterodyne spectroscopy, ground-based microwave radiometry, 

and limb scanning at visible and microwave frequencies are noted. 

ANALYSIS OF OZONE VARIATIONS 

DATA AVAILABILITY 

Ground-Based Observations 

As previously mentioned, ground-based total-ozone observations are currently made in a global network employing Dobson ozone spectrophotometers (at about 53 stations), and broadband filter ozonemeters 

(at about 22 stations). Data from the stations are routinely published by 

the Canadian Atmospheric Environment Service, Department of the 

Environment (Ozone Data for the World (ODW)), in cooperation with 

the World Meteorological Organization. Figure 2 is a plot of the global 

distribution of total-ozone observing stations with an indication of those 

in current operation. As can be seen, most of the stations are at midlatitudes in the Northern Hemisphere and on land. 

A few Umkehr observations were made prior to the start of the International Geophysical Year (IGY). However, quasi-regular data have been 

available only since 1958. The present network consists of four stations 

in Europe, four in Japan, five in India, and four in Australia. Any 

Dobson station has the capability of making Umkehr measurements. 

Balloons 

Although many balloon data have been accumulated by limited flight 

programs and published independently, a few synoptic networks have 

been operated. The U.S. Air Force Cambridge Research Laboratory 

derived synoptic results from a North American network from 1964 to 

1966. At about the same time, the Environmental Sciences Service Administration (now known as the National Oceanic and Atmospheric Ad-. 

ministration) established a meridional array of sounding stations in both 

hemispheres. An international European network of four stations was 

established in the late 1960's. Long-term soundings have been made at a 

few stations (e.g., Thalwil, Switzerland; Weissenau, Hohenpeissenberg, 

and Berlin, Germany; Wallops Island, Virginia; and AspendaIe, Australia). 

The results from the networks and time series are generally available as 

in-house documents and special reports. The more recent results are 

published in Ozone Data for the World. 
68 CHLOROFLUOROMETHANE WORKSHOP REPORT 

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The number of rocket soundings ismuch smaller than the number of 

balloon soundings because rocket flights are genray made for speciai 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 69 

experiments rather than for climatological purposes. The cooperative 

Meteorological Rocket Network is considering a program of routine 

flights, and the National Aeronautics and Space Administration (NASA) 

Existing is conducting a series of monthly soundings from a few stations. 

rocket data are available in individual publications, and Krueger and 

Minzner (1976) have constructed a "standard 45°N ozone atmosphere" 

from these data. 

The distribution of routine stratospheric ozone measurements (Umkehr, 

ozonesonde, and rocketsonde) both past and present is shown in figure 3., 

Again it can be seen that most of the stratospheric ozone observations are 

currently being made over land areas in midlatitudes of the Northern 

Hemisphere. 

•W
 

;-1 II Z I.-

LEGEND * UMKEH'R MEASUREMENTS (DOBSON) 

A 0Z0NESONDE (ELEC'rROCHEMICAL.CONCENTRA'TlONCEELL
 

U ROCKETSTONDE
 

UNDERINE IND[CAESECU RRENT OPERATION ASOF 1975.
 

Figure 3. Observations of stratospheric ozone (1957-1 975). 
70 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Satellites 

By far, the most extensive satellite data set currently in existence is that 

obtained with the BUV experiment carried on Nimbus-4 (Heath et al., 

1973). Data have been taken over approximately the past 7 years (since 

April 1970) and are still being acquired on a limited basis. Both highlevel profiles and total-column amounts can be retrieved from these data. 

The latitude coverage is 80°S to S0 N, and all data are acquired at 

approximately local noon. Although only small samples have been reduced and published thus far, an intensive effort is currently underway 

to process the entire body of data and make it available to the scientific 

community. This task is scheduled for completion by mid-1978. 

A spare Nimbus-4 BUV instrument was flown on the AE-E satellite that 

was launched in December 1975. Again, total-ozone and upper-level profiles are obtained by this instrument. Although the latitude range 

covered is only 200S and 20'N, more nearly complete coverages in 

longitude and local time are possible than could be achieved with the 

Nimbus orbit. Preliminary reduction of limited samples of data have 

thus far been carried out. 

Observations made on the polar orbiting satellite, OGOA, have provided 

data for calculating the global distributions of the vertical ozone profiles 

in the range 10 millibars to 0.4 millibar. Measurements are available for 

the period September 1967 to January 1969. The distributions for the 

period September 1967 to February 1968 have been calculated (London 

et al., 1977), and the remaining data are now being processed. 

The Nimbus-6 LRIR experiment (Gille 	et al., 1975) has provided limb­ " radiance measurements in the 1042-cm band over a 7-month period 

from June 14, 1975 to January 7, 1976 (when the subliming cryogenic 

material was exhausted as anticipated). These data cover the globe from 

approximately 80'N to 65°S, with most daytime observations at approximately 1330 local time and night observations at about 2230 local time. 

Preliminary profiles have been retrieved from these data, and a finalized 

data set is expected by mid-1977. 

More than 3 months of data (April through July 1969) were obtained 

from the Nimbus-3 IRIS, and approximately 9 months (April through 

December 1970) of total-ozone estimates were obtained from the 

Ninbus-4 IRIS experiment. These data, taken either at local noon or 

local midnight and covering the latitude range from 80WS to 80N, have 

been described in detail by Prabhakara et al (1976). 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 71 

Future Satellite Data 

In early 1977, a meteorological sounder on a Block 5 satellite began 

returning measurements of, among other things, upwelling radiation in 

the 1042-cme r band. This series is planned to maintain two satellites in 

orbit into the 1980's. 

The Nimbus-G research satellite, to be launched in 1978, will carry 

improved instrumentation for obtaining both backscattered ultraviolet 

measurements and horizon-emission measurements of ozone. 

The SAGE will measure high-resolution ozone profiles by observing solar 

occultations in the Chappuis band. A 1978 launch is planned for an 

Applications Explorer Mission (AEM) satellite. 

It should be noted that there are currently no long-term plans to monitor 

either the ozone profile or the total ozone by satellites. 

LONG-TERM VARIATIONS 

Total Ozone 

Ground-Based Data 

In discussing total-ozone trends derived from the Dobson instruments, 

two basic limitations of these instruments for estimating hemispheric or 

global variations should be emphasized. First, the instruments may go 

badly out of calibration every few years, and the lack of calibration has 

been a continuing problem with this network. Second, the distribution 

of Dobson stations is not adequate for accurately estimating hemispheric 

or global variations, mainly because the stations were established regionally (mostly during the IGY) for studying the synoptic variability in 

total ozone. The problem of using a heterogeneous network has been 

compounded by frequent activation and deactivation of Dobson stations, 

and it is therefore difficult to estimate the representativeness and reliability of these data for determining total-ozone trends. The best Dobson 

network is to be found in the north-temperate latitudes, in which there 

are two regions with good Dobson coverage (Europe and North America) 

and two regions with adequate coverage (Japan and India). 

Investigators using different methods of statistical analysis have inferred 

an increase in total ozone during the 1960's and generally some decline 
72 CHLOROFLUOROMETHANE WORKSHOP REPORT 

following 1970. Komhyr et al. (1971), using a regression analysis, showed 

that total ozone increased by about 5 percent at most North American 

stations during the 1960's. Although it was initially considered possible that this was strictly a regional effect due to a shift in the longwave pressure pattern (the total-ozone amount tends to be higher in 

troughs than in ridges), London and Kelley (1974), using grid-point data 

from mean monthly maps, showed that an increase of this magnitude 

occurred over much of the Northern Hemisphere during this period. 

Goldsmith et al. (1973), Christie (1973), and Johnston et al. (1973), 

using diverse analysis techniques, also detected an increase of total ozone 

during this period. Figure 4 shows the total-ozone variation (expressed 

as a percentage deviation from the record mean) in climatic zones, based 

on the analysis of Angell and Korshover (1976). 

12 1 I 1 1 1 

4z
 

o
 

.4 20


4
 

No A. --2
 

0
 

4 

00 

M-2 


z-­

z-, 4 w 

0 

1950 1955 $960 1965 197/0 1976 1980 

Figure,4. Time variation in total ozone for various climatic zones. Variation is expressed as a percentage deviation from the mean for the length 

of record. The annual oscillation has been eliminated and a 1-2-1 smoothing applied to the successive seasonal values. The vertical bars extend two 

standard'deviations of the mean on either side of the mean, whereasthe 

number of Dobson stations is indicated in parentheses at right. 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 73 

Note that the annual oscillation has been eliminated by taking deviations 

from the long-term seasonal means and that the vertical bars embrace two 

standard deviations from the mean. If the vertical bars do not overlap, 

the total-ozone variation is assumed to be significant at the 5-percent 

level. In north-temperate latitudes, the total-ozone amount is seen to 

have increased by about 5 percent between 1963 and 1970, and, since 

J970, it has increased by about half this amount. This variation is 

mostly the result of rather large variations in Europe and North America, 

with only a slight long-term trend in Japan and India after 1960. In the 

USSR, the increase in total ozone during the 1960's was about four 

times as great as that indicated by European and North American stations, and, accordingly, the Soviet filter ozonemeter data were not included in the north-temperate latitude average. The single north-polar 

station' (Resolute, Canada) exhibits a, trend similar to that found for 

north-temperate latitudes. 

The Dobson network in the tropics is rather sparse and now consists of 

only Huancayo, Peru (in need of calibration); Mauna Loa, Hawaii (more 

a subtropical than tropical station); and Kodaikanal, India. The inadequacy of the tropical Dobson network is especially unfortunate because, 

if CFM does affect the ozone, the effect should be first noted near the 

Equator where the synoptic and annual variations in total ozone are 

small, where the photochemical effects dominate the atmospheric transport effects, and where CFM's are apparently transported upwards 

through the tropical tropopause by means of the Hadley cell circulation 

(toward the region above 30 kilometers in which the CFM's are dissocia-, 

ted by ultraviolet radiation). Establishment of an adequate Dobson network in the tropics is of utmost priority, as indicated by recent WMO 

recommendations. In the meantime, the total-ozone variation in the 

tropics illustrated in figure 4 should be accepted with great caution, as 

emphasized by the lengths of the vertical bars. 

In south-temperate latitudes (Australian stations and Buenos Aires) and 

south-polar latitudes (Amundsen-Scott or South Pole), the total-ozone 

trend has been slightly downward during the period of record. Superimposed on this trend is a pronounced quasi-biennial oscillation in total 

ozone at the Australian stations. It is obvious that the Dobson network 

is also inadequate in south-temperate and south-polar latitudes, and thefe 

are recent WMO recommendations in this respect. 

Figure 5 shows the derived variation in global total ozone obtained by 

weighting the climatic zones (figure 4) by the area of the Earth's surface 
WORLD 

4 

o 

C 

-2 

N0 

-4 

-5 

1960 1965 1970 

Figure 5. Time variation in total ozone for the world. Variation is obtained by weighting 

the climatic zones by the approximate area of the Earth's surface that they represent. 

1975 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 75 

that they represent. The resulting global trends move significantly upward during the 1960's, move significantly downward during the early 

1970's, and tend to zero during the mid-1970's. The slight (and provisional) increase indicated for 1976 results from the dominating effect 

(because of the relatively large area involved) of the tropical trend; as 

previously mentioned, the sparseness of data 'in this region make the 

analysis somewhat doubtful. 

It is of interest to compare the global total-ozone trend in figure 5 with 

the trend obtained by London (unpublished) by averaging individual station values in latitude bands and then weighting these bands by the area 

of the Earth's surface that they represent. London obtained about a 

5-percent increase in global total ozone between 1963 and 1970 in comparison with the 3-percent increase noted in figure 5 and a decrease between 1970 and 1973 of about 1.5 percent in comparison with the 2­

percent decrease indicated in figure 5. This illustrates the relatively small 

difference in total-ozone trends obtained by applying different analysis 

techniques to the same data. A noteworthy feature of London's results 

is the increase in trend magnitude with increasing latitude, as well as the 

increase in shorter term variability with increasing latitude. 

A factor to be considered in deducing total-ozone trends from data obtained at stations located in industrial regions of the globe is the effect 

of anthropogenic influences on the measurements. White et al. (1976) 

have reported that ozone concentrations in excess of 20 pphmv have 

been observed in urban pollution plumes that extend several hundred 

kilometers. Since background suiface-ozone concentrations in clean air 

amount to 3 pphmv, the bulk of the observed surface ozone must be 

produced photochemically within the plume. Assuming a mean vertical 

height for a plume of 1 kilometer, a concentration of 20 pphmv of 

anthropogenically produced ozone amounts to an ozone contribution to 

the total atmospheric ozone column of 20 m-atm-cm (or 6 percent of the 

ozone for a total-ozone amount of about 300 m-atm-cm). At Dobson 

spectrophotometer stations located in areas subject to pollution, therefore, it appears worthwhile to make measurements of surface ozone along 

with observations of total ozone in order'to estimate the contribution of 

photochemically produced surface-ozone variations to variations in background total-ozone amounts. 

Satellite Data 

As mentioned in the section on availability of satellite data, only a small 

amount of satellite-derived total-ozone data have been reduced, and published. At the present time, comparison of absolute concentrations of 
76 CHLOROFLUOROMETHANE WORKSHOP REPORT 

ozone determined from satellites of nonoverlapping time periods is not 

particularly useful for ascertaining trends, because there can be a larger 

uncertainty in the absolute calibrations than the likely amplitudes of the 

trends. The potentially most useful of the present data sets, that from 

Nimbus-4 BUV, has been reduced for the first 2 years of operation. For 

this period, the total-ozone measurement indicates a decrease of about 

0.30 percent, which is in qualitative agreement with the trend deduced 

from the ground-based data (see figure 4). Intercomparison of results 

from Nimbus-4 with those from different satellites at overlapping times 

(e.g., AE-E) and with the Dobson network will lend confidence to the 

absolute values over the period through 1976. 

Stratospheric Ozone 

Umkehr Data 

Relative to measurements of total-column abundance of ozone, stratospheric measurements of ozone density are probably of greater importance in early detection of ozone depletions attributable to CFM's. It is 

shown in Chapter 4 that the reduction in ozone attributable to the effect 

of CFM's would preferentially occur near the region where the CFM's are 

photodissociated, and, in this region (35 to 45 kilometers), the percentage 

ozone reduction would be almost nine times the reduction observed in 

total ozone. The satellite BUV and limb-scanning measurements should 

be excellent for this purpose in the future, but somewhat less precise data 

are now available from the Dobson instruments using the Umkehr 

method. 

Figure 6 shows the ozone trends in the layers 32 to 37 kilometers, 37 to 

42 kilometers, and 42 to 47 kilometers for the four regions (Europe, 

Japan, India, and Australia) over which Umkehr measurements have been 

obtained. The following can be seen: 

* 	 The ozone trends at these high levels are similar to the trends in 

total ozone. (Some similarity might be expected a priori because 

the Umkeh measurements are forced to sum to the Dobson measurements of total ozone.) 

* 	 The amplitudes of the ozone trends appear to increase with height. 

It is uncertain whether this variation with height is real or is somehow inherent in the Umkehr technique. (Rocketsonde data do not 

suggest such an increase with height.) 
12 

1I 32-37 lru I-''I 37-42 r km - rI I 42-47 kmI 

-00 

-4 SZ 

4z 

-4 

-8-

M 

Ao 

-8 

24 I 

-12 

20 

-28 

1qIIII ej 

0 

0C 

Z 

1960 1965 1970 1957 1960 1065 1970 1975 1960 1965 1970 1976 

Figure B. Time variation in ozone amount at heights of 32 to 37 kilometers (left), 37 to 42 kilometers 

(middle), and 42 to 47 kilometers (right) for various regions (from Dobson Umkehr measurements). 
78 CHLOROFLUOROMETHANE WORKSHOP REPORT 

" 	 A large ozone depletion was indicated by the two Australian stations at heights above 40 kilometers following the eruption of Mt. 

Agung in 1963. A more detailed analysis shows that the decrease 

occurred later in time the further away in latitude the stations were 

from the eruption site, as if a haze cloud were dispersing poleward 

from the volcano. Because the volcanic debris could hardly have 

exceeded 30 kilometers in height but the implied effect on ozone 

increases up to 45 to 50 kilometers, it is likely that the indicated 

ozone depletion is fictitious (or mostly fictitious) and attributable 

to the effect of volcanic aerosols on the Umkehr measurement 

technique. 

* 	Between 1974 and 1975, a rather abrupt decrease in ozone occufred at the higher levels above Europe and Japan, somewhat 

similar to the total-ozone decrease noted in North America in 1976. 

It is of interest to compare these Umkehr trends with the trends observed 

from ozonesonde measurements in Switzerland (London and Dutsch, 

1976). They show that the increase in total ozone between 1966 and 

1970 in this area was associated with an increase in ozone in the layer at 

20 to 25 kilometers, but that there was very little trend at higher levels. 

This is somewhat inconsistent with the Umkehr results for Arosa, Switzerland, as discussed by Dfitsch and Ling (1973), and comparison of the 

results of these two methods needs to be pursued further. 

Balloon and Rocket Data 

Flight platforms available for sounding this region are balloons, rockets, 

and satellites. The ozone profile from 2 to 74 kilometers shown in figure 7 

is a composite of rocket and balloon data from the midlatitudes (Krueger 

and. Minzner, 1976). The rocket data were taken over a period of 7 years 

by diverse techniques and thus give an upper limit to the variability. The 

mean variability between 30 and 50 kilometers is about 15 percent, in 

contrast to values approaching 100 percent below 20 kilometers. The 

low variability near 40 kilometers, combined with the greater response of 

ozone to CFM's at these levels, suggests an improved signal-to-noise ratio 

for detection of changes in this region. 

SatelliteData 

Measurements from the LRJR and BUV that have not yet been analyzed 

will give many more determinations of ozone in this region, from which 
80 

70 

Cn 

50 

40 

30 

tu 

TH 

TO 

10­

21014 1015 10I 

OZONE CONCENTRATION 

1017 

(molecule~im 3 1 

101i 1019 

0 

Figure 7. Model of the midlatitude vertical ozone distribution and its variability based on 

balloon and rocket data. The variability indicated by horizontal bars shows lower values 

between 30 and 50 kilometers than at other altitudes. 
80 CHLOROFLUoROMETHANE WORKSHOP REPORT 

seasonal variation may be determined and the unexplained variability 

reduced. Existing measurement programs can only marginally detect predicted changes in stratospheric ozone at this time, as in the case of total 

ozone. However, because increased depletion is expected and because of 

improved satellite and rocket programs, a conclusive measurement can be 

expected by 1980. An improved BUV instrument and a slightly modified 

LRIR (denoted Limb Infrared Monitor of the Stratosphere (LIMS)) will be 

flown on Nimbus-G in 1978. Rockets are being systematically launched 

to bridge the data gap between Nimbus-6, AE-E, and Nimbus-G. This 

coordinated effort is capable of early detection of changes. 

SOME INFERENCES ON PHOTOCHEMICAL MODELS DERIVED 

FROM OZONE-VARIATION OBSERVATIONS 

Observations of ozone variations in the middle and upper stratosphere, 

in which photochemistry is dominant in determining the ozone concentration, can be used to further the understanding of the ozone chemistry 

in this" region. The following subsections describe two examples (of 

many) in which such comparisons have been particularly useful. 

Relationship Between Ozone Concentration and Temperature 

By combining temperature observations and BUV ozone concentrations 

measured from the Nimbus satellite during a stratospheric warming at 

levels above 2 millibars, Barnett et al. (1975) found an inverse correlation between temperature and ozone concentration. Using simple photochemical theory, they showed that oxygen (and odd nitrogen) chemistry 

gave a much larger temperature dependence than odd hydrogen chemistry. Comparison between the calculated and observed temperature dependence showed values between the pure oxygen and odd-hydrogen 

chemistry values. This could indicate that the odd-hydrogen chemistry 

does not operate as effectively as believed or that the temperature dependence of the rate coefficients may not-be correct. 

Solar-Proton-Events 

The effects of continuous injections of ozone catalysts in the stratosphere should be detectable as long-term changes in total ozone or its 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 81 

vertical distribution. These slow trends tend to be obscured by other 

sources of ozone variability, such as changes in atmospheric circulation. 

However, transient injections may be uniquely'identified by their spatial 

and temporal signatures. In early August 1972, a sudden decrease in the 

density of upper stratosphere ozone at high geomagnetic latitudes was 

observed with the Nimbus-4 BUV (Heath et al., 1977) (figure 8). This 

change was coincident with a major solar-proton event, which should 

have enhanced the nitric-oxide concentration. The altitude distribution 

and amplitude of the ozone decrease were substantially as predicted in 

calculations by Crutzen et al, (1975). Thus, the catalytic effect of nitric 

oxide within the stratosphere was quantitively supported by this observation. 

0015 75-80N 

Cc 

0013 +~+ 

0 

m 

N
 4I+--14W+HI 

0-003
 - i 1 1 llll1 1li lil/ ll lli lll ~f~l~lll11l1.,i 1 

1 10 20 1 ,10 20 1


JULY AUG SEPT 1972 

Figure 8. Decrease in the zonally averaged integral amount of ozone 

above the 4-millibar surface for the 75 to 80ON zone associated with a 

solar-proton event on August 4, 1972, as observed with the Nimbus4 

BUV. The sudden decrease in total ozone in early August 1972 was 

coincident with a major solar-proton event. 
82 CHLOROFLUOROMETHANE WORKSHOP REPORT 

COMMENTS ON THE ANALYSIS OF OZONE DATA 

DISTINGUISHING BETWEEN NATURAL AND 

ANTHROPOGENIC TRENDS 

A major concern involving determination of CFM-produced ozone variations deals with the threshold of detectability, if indeed there is (and/or 

has been) a significant effect on the ozone concentration of past, present, 

and future CFM releases. 

The problem of the threshold of detectability is made particularly difficult because of the high natural variability of the observed data. As 

pointed out in the foregoing discussion, ozone measurements exhibit 

many periodic variations (e.g., seasonal and quasi-biennial) and are 

strongly related to meteorological variabilities that tend to mask or obscure long-period ozone changes. Thus, the early effects of any anthropogenic perturbation could be so small as to be indistinguishable from 

the natural variations. Statistical techniques need to be used that will 

minimize the difficulty of distinguishing between natural and anthropogenic variations. In addition, it is necessary to determine if changes that 

are observed in the ozone concentration are global in extent or simply the 

result of a redistribution of ozone due to meteorological variations (e.g., 

shifting of atmospheric circulation patterns). 

Total ozone* at a particular location can be considered a random or 

stochastic function of time, 03 (t), made up of two distinct components: 

periodic oscillations and a nonperiodic component. The observed time 

variation of ozone is then given by: 

0 3 (t) = 03 + 2 Fi 0 3 (t) +X'(t) 1 

where 

0 3 = the long-term average 

Fi 03 (t) = the ith periodic component 

X' (t) = the residual nonperiodic component 

*Reference to total ozone could be just as well applied to ozone concentration at 

some level. 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 83 

The periodic part can be modeled as the sum or superposition of several 

periodic functions of time, each with a different amplitude, phase, and 

period. These amplitudes, phases, and periods may be random processes 

themselves so that the periodic variations may be called "quasi-periodic." 

Without loss of generality, it can be assumed that the time average (averaged over all time) of the periodic component of 03 (t) iszero. 

The nonperiodic component can then be reduced to a function of time, 

X(t), plus white noise, E(t), having a variance, 0 (t). For the purposes 

of discussion, assume that 4 (t) = 4, (i.e., the variance of the white 

noise does not change with time). Thus, the nonperiodic component of 

03 (t) might look like the graph in figure 9. 

X (dl
 

Figure 9. Time variations of the nonperiodic 

component of ozone (schematic). 

X'(t) can be filtered by various standard methods to give X(t) a smooth 

function of time. The total ozone will be said to have no trend if X(t) = 

constant (that is, if X(t) is stationary). A variation of X(t) with time will 

be a trend. In this definition, nothing is said about the cause of the 

trend-whether it is "natural" or a result of man.'s intervention-or how 

long it persists. 

Note that the definition of the periodic variations as having a zero time 

average presupposes a knowledge of 0 3 (t) for all time. In practice, only 

some finite interval of time, T, can be observed. In the case of total 

ozone, the largest value of T, that at Arosa, Switzerland, is less .than 50 

years, whereas most observing stations have less than 20 years of data. 

This implies that some natural oscillatory waves may have periods that 

cannot be distinguished from nonperiodic trends. Thus, apparent trends 

in a data set can, in fact, be parts of long-wave natural oscillatory variation. 
84 CHLOROFLUOROMETHANE WORKSHOP REPORT 

STATISTICAL ANALYSIS APPLIED TO PRESENT DATA SETS 

Virtually every statistical analysis seeking to detect a trend attempts in 

one way or another to remove or "filter out" the oscillatory component 

of the data and to isolate the nonperiodic component. 

Three basic steps are usually taken. The first involves data smoothing in 

which "raw" data are replaced by averages of data that are close together 

in time and space. This smoothing removes some of the high-frequency 

variations and some of the white noise. The next step is to determine the 

different periodic variations. Any oscillatory variation that is not removed in this process remains as a part of the "residual." The third step 

is an examination or modeling of these residuals in an effort to estimate 

X'(t) and X(t). The consequence of leaving unevaluated periodicities in 

the residuals is that the estimate of Sx (the standard error of the estimate) becomes inflated. Since the errors made in estimating X(t) are proportional to Sx, a goal of any trend analysis should be to remove as much 

significant periodic variations as possible to reduce the variance in the 

residuals. 

A method that reportedly reduces Sx significantly is the time-series analysis discussed by Hill and Sheldon (1975). According to their analysis, 

residuals in the total-ozone data from Arosa, Switzerland, appear to be 

- white noise, and there is no significant trend for the entire data period 

(1932 to 1974). It should be noted, however, that a significant upward 

trend of total ozone was noted by Komhyr et al. (1971) in the data for 

Arosa for the period 1960 through 1970. As Birrer (1974) has shown, 

trends in past Arosa data lose their significance with an increase in the 

period for which the trends are calculated. A model of the Arosa data 

prior to 1970 was used by Hill and Sheldon (1975) to predict the observed downward trend through 1974. This prediction is largely a consequence of the continued propagation of the unexplained 133-month 

period in the ozone record at Arosa as described by them. 

THRESHOLD OF DETECTABILITY 

Errors made in estimating the slope, 3,of a linear trend are proportional 

to Sx/Fn, where n is the number of independent observations. For any 

Sx, the minimum detectable slope can be determined as a function of n. 

Since this "minimum detectable trend" is also dependent on Sx, the 

result is dependent on the statistical method used to remove the periodic. 

components of 0 3 (t). 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 85 

Using their model involving data from a sample of. nine globally distributed stations, Hill et a. (1977) estimated that a trend producing a 

1.56-percent change over 6 years would be statistically significant. This 

means that if a linear trend with slope 0 = 0.26 percent per year begins at 

timeto, then,by time = 6 years, this linear trend will be recognizable from 

the zero trend observed in the larger data set prior to time t o (shown in 

figure 10)_ 

F.ct taf 

Ile 

pfa 

.... men d pt 

base 10. howellmati repesnttion of thereidua stratistic of the large 

data set prior to time t) can be reduced. The level of detectability 

claimed by Hill et al. (1977) is lower than that previously estimated by 

Pittock (1974) for a global network. A large part of the difference in the 

two estimates is probable due to the fact that the statistical analysis used 

by Hill et al. (1977) resulted in smaller residual variances. Pittock did not 

remove the variance due to the mean annual cycle from his data. 

If a "statistically significant" trend is observed, it could indicate that X(t) 

is changing due to some disruption in the equilibrium that existed prior 

to time t, or that the model used to remove periodic components is 

breaking down due to shifts in periods, amplitudes, or phases that may 

themselves be random. 
86 CHLOROFLUOROMETHANE WORKSHOP REPORT 

SUMMARY AND CONCLUSIONS 

The present estimate of the effect of CFM's injected into the stratosphere is a long-term reduction of total-column abundance of ozone of 

about 7 percent with a maximum reduction of the order of 60 percent 

in ozone concentration in the middle and upper stratosphere (35 to 45 

kilometers). A careful analysis of the past records of total ozone and its 

vertical distribution contributes to an understanding of the origin of natural variations of ozone (both periodic and nonperiodic) and therefore 

should aid in detecting ozone variations that are related to the predicted 

reduction. 

The data base for determining perturbations in the ozone series is best 

documented in the observations of the total amount. However, because 

the largest percentage of reduction in ozone is expected to occur 

at about 40 kilometers and to occur earlier at this level than in the 

total amount, special efforts should be made to provide information 

on the natural variations at this height and to detect the predicted CFM 

effects. 

Many measurement techniques are available for determining total ozone 

with the use of ground-based and/or satellite platforms. An obvious and 

important advantage of satellite observations is the potential for global 

coverage, and continued satellite total-ozone monitoring programs need 

to be assured. The ground-based measurements have the benefit of a long 

history of observations with about 75 observing stations operating in the 

present international network, some of which have taken observations 

more or less continuously for 20 years or more. Although the precision 

of these techniques varies somewhat, it is of the order of ±3 to +5 percent 

for ground-based methods and about ±5 percent for satellite determinations of total ozone. It should be emphasized that a coordinated program 

of complementary -ground-based and satellite measurements is capable of 

providing better quality determinations of total ozone than those that 

can be obtained with either technique alone. 

The best methods for determining long-term trends in the vertical ozone 

distribution are through Jmkehr observations because they provide a 

relatively long-term data base or through the use of balloonborne and 

rocketborne or satellite platforms, with the latter two the more promising 

for the region above 30 kilometers. In addition, rocket observations 

present an excellent way to calibrate the satellites, and a coordinated program of in situ and satellite measurements holds the best promise for obtaining good-quality profile information. 
MEASUREMENTS AND ANALYSIS OF OZONE VARIATIONS 87 

The variability and long-term trends in total ozone are closely related to 

the variability and long-term trends at and just below the level of ozone 

maximum. Indications are that the natural variability of the ozone concentration decreases above the ozone maximum in the tropics and midlatitudes. At high latitudes during winter, however, the variability is quite 

large with evidence of large fluctuations associated with meteorological 

perturbations (e.g., sudden warmings). Thus, observing systems designed 

for determining the presence of CF'M-induced trends in the middle stratosphere should be planned to focus on those atmospheric regions in which 

the natural periodic and nonperiodic variations are smallest and in which 

photochemistry is the dominant process that produces ozone variations. 

Analysis of past total-ozone data indicate the broad features of the temporal ozone changes (i.e., seasonal variations), as well as the general increase in ozone in the Northern Hemisphere during the 19 60's. However, 

the seasonal variations and long-term trends in the stratosphere, particularly above the level of maximum ozone are not nearly as well-documented. Extension of the ozonesonde network would contribute important information on ozone variability in regions of the largest concentrations. However, it is essential that provisions for a long-term stable satellite and rocketsonde program be made to fill the gap of ozone information in the stratosphere. 

Although presently available statistical techniques appear to be adequate 

for filtering the various "natural" periodicities in the ozone records and 

for giving estimates of the variance of the residual data, work now in 

progress gives promise that statistical analysis of long-term ozone trends 

can be improved to give more definitive results of possible interruptions 

(or modulation) of the present patterns by CFM inputs to the atmosphere. However, it is clear that a larger data set (based on more precise 

observational techniques) than is currently available is needed to verify 

the rate of ozone destruction by CFM's postulated by the models discussed in this document. 

Regarded as being generally correct and timely is that part of the general 

statement contained in the report of the Committee on Impact of Stratospheric Change (Tukey Report), which states that studies must be conducted dealing with: 

"... Development and general acceptance of standard methods for 

routine mapping and summarization of the measurements of the 

distribution in space and time of ozone and other low-concentration constituents. The reduction of such measurements is far from 
88 CILOROFLUOROMEThANE WORKSHOP REPORT 

trivial and requires well-chosen allowances for fluctuations both in 

smoothing and in stating the uncertainties ofthe results." 

* 	 "... Processes and mechanisms determining the natural variations 

of ozone, particularly the longer-term changes. As we improve our 

understanding of natural ozone changes, we will be able much more 

easily to interpret observed changes, to identify human-produced 

changes, and to give guidance to the assessment of historical 

changes in UV-B intensity." 
N77 -316 97
 

CHAPTER 3 

MINOR SPECIES AND AEROSOLS 

INTRODUCTION 

In the present state of the art, stratospheric measurements are available 

only for some of the most critical species. In some cases, successful 

techniques have not yet been developed. Nevertheless, a variety of species 

have been measured and the existence of critical species in the stratosphere 

has been established. Ih this chapter, an attempt is made to establish some 

order to the measurements that have been made and in the type of 

problems that exist. 

A measurement of the species that represent sources and sinks as, a function of latitude can be of considerable help in tying together the efforts of 

one-dimensional (1-D) and multidimensional models. -Given appropriate 

information off latitude variations, boundary-value terms for the equations 

can -be developed. In this case, efforts to enhance the precision and crosscorrelation of results among various investigators are important. 

In the evaluation of sinks and sources in the troposphere, the limiting 

factor is the accuracy with which measurements can be made. At present, 

there are no standards by which.various groups can verify the absolute 

accuracy of their techniques. The development of such standards is necessary if accurate estimates of tropospheric sinks are to be possible. 

Another type of constituent measurement is that which, checks models 

involving the creation and destruction of radicals. For these models, 

measurements with high accuracy are not as important as good precision 

and the availability of a variety of correlated measurements. Each of 

these requirements will be discussed later in this chapter. 

The physical and chemical processes that occur in the stratosphere are so 

complex and coupled that measurements of the chemical species and 

physical parameters involved are necessary for obtaining an understanding 

of this region of the atmosphere. These measurements can, and do, provide the knowledge required for identifying individual phenomenon, for 

assessing their importance, and for providing-necessaiy quantitative checks 

and constraints on theoretical descriptions of the processes involved. 

These checks assume increased importance when theoretical descriptions 

are used for prognostication, as in the case of the evaluation of chlorofluoromethane (CFM) influence on stratospheric ozone. 

89 
90 CHLOROFLUOROMETHANE WORKSHOP REPORT 

The ultimate interest in this assessment is whether total-ozone column 

density and ozone distribution is changed by the release of CFM's into 

the lower atmosphere. This interest necessarily involves the identification and consideration of the processes responsible for ozone loss and 

production. It involves a scale that is global, evaluation of photochemical 

processes involving species that exist in the stratosphere, and interaction 

of the production and loss of trace species with the general circulation 

and other transport processes that occur in the atmosphere. The state of 

empirical and theoretical knowledge concerning many of these species 

and processes is at a rudimentary level. 

Many species are important to the calculation of ozone destruction. 

Those species that have been identified as being most important are listed 

in table 25 along with pertinent data concerning the extent to which they 

have been measured. In addition, a variety of ratios are important in the 

chemistry of ozone. In some instances, these ratios can be shown to be 

determinable from relatively simple measurements. Chief among them 

are the 0/03 ratio, the NO/NO2 ratio, and perhaps the CI/CIO ratio. In 

almost all instances, these ratios depend strongly on the solar-flux value 

that dissociates the principal molecule. The solar flux from about 180 to 

600 nanometers is the region of most importance for a variety ofstratospheric species. It is well known that the flux in the region below about 

130 nanometers varies with solar activity. It is also known that the solar 

constant appears to be constant to a few percent when measured from the 

ground. The spectral region between about 150 and 350 nanometers is 

not easily monitored, and less is known about the constancy of the Sun in 

this region. In the past, it was generally assumed that the critical region 

between 200 and 300 nanometers did not vary in time. Recent results 

indicate that this assumption may be in error (Heath and Thekaekara, 

1976). At present, no solid experimental data exist on which either the 

constancy or the variability of the Sun in this spectral region can be 

determined. 

A parameter that has only recently been considered is the error introduced by assuming single scattering models (Callis et al., 1975; Luther 

and Gelinas, 1976; and Luther and Wuebbles, 1977). At present, few experimental data exist that can be used to establish the validity of single 

versus multiple scattering models. This is an area in which experimentation is needed. 

In the discussion of photochemistry, a phenomenon often referred to is 

mixing or transport. The effects of these phenomena and the ways in 

which they are parameterized in models are extensively discussed in other 
Table 25 

Important Species 

Specits Aes 

tNoof 

Mra. 

Teeh 

nique' 

Ilight 

Profile 

Liliude 

Ptolc 

SeaSonal 

Varlstlo.n 

Temporal 

Vaiatuns 

Ilemisphenea 

Variation 

(FCI 

CFCI 

(C14 

NO 

Cll 

Yes 

Yes 

Yes 

Yes 

Yes 

Many 

Many 

Many 

.10 

> 10 

a 

a 

a 

a,b,d 

a,1, 

035 

043 

021 

55 

1250 

Yet 

Yes 

Yet 

-

-

Yes 

Yes 

Yes 

0 

NO 

NO2 

[i2O, 

Cl 

I0 

Yes 

Yes 

Yes 

Yet 

Yes 

Many 

-10 

-5 

2 

3 

cb 

h,b 

d, b 

g 

h,,g 

17-60 

1037 

12f3 

3544 

2342 

Yet 

Yo, 

Yes 

-­

-

Yes 

Yes 

? 

Yes 

Yes 

Y's 

-

CO 

C13 

CIONO, 

lCI 

I10 

No 

Yes 

Yes 

Yes 

2 

-l0 

> 6 

b 

d, b 

g.i 

No 

1232 

070 

--

YOs -'r 

Yes, colun 

tv 

1102 

UlO 

NO 

No 

abundft¢c Go

0 

1120 Yes, .ny C, 1',1, 

llr 

sir 

Yes 

Yes -4 

d 

b,d 

Yes 

No Yes No 

0 

1i 

Yes 

Yes 

2 

7 

6 

a 

45.27 

150 -

-, 

Total 

didlne No 

Ii - Inned *p~t~t ecrwi yg urlaetlccnccl e asmui'S nCo enrahlos C hsyersmecig hI vkin atnleultsa'0ehItabosplen sp¢nn I * tlueaucr . 
92 CHLOROFLUOROMETHANE WORKSHOP REPORT 

chapters of this report. In this chapter the few measurements that are now 

available are discussed. 

In considering the mixing processes in the stratosphere, it should be noted 

that most height and many horizontal profiles of trace constituents 

present large variability over apparently very small scales. The term 

"'transport" is usually invoked to explain these variations. Mechanisms 

are rarely discussed, however, by which processes that move air can alter 

the concentration of a trace constituent. If such mechanisms do not 

exist, then air must have been moved in from another altitude at which 

the observed concentration is prevalent. At present, few data sets exist 

that are complete enough to permit an accurate description of such 

processes, especially on a mesoscale. 

SUMMARY OF ATMOSPHERIC MEASUREMENTS 

The strong chemical link among the chlorine, nitrogen, hydrogen, and 

oxygen photochemical systems in the stratosphere dictates that discussions of the field measurements be divided not according to chemical 

identity, but rather according to: (1) whether the species constitutes a 

chemical source that in general enters the stratosphere through the tropopause, is destroyed by reaction or photolysis, and is not, reformed in a 

further photochemical step; (2) whether the species is a chemical sink 

or "reservoir" term that is a product of the photochemical reactions 

taking place in the stratosphere and that is carried downward out of the 

stratosphere into the troposphere; or (3) whether the species is a reactive 

fragment or radical that is produced from the source term by photolysis 

or oxidation and that recombines to form the sink terms. Advancements 

since the National Academy of Sciences (NAS) report in each of these 

areas are discussed in the following sections. 

SOURCE SPECIES 

Table 26 presents an inventory of the tropospheric mixing ratio of molecules believed to be of importance to the photochemical budget of the 

stratosphere. Particular attention is given to the question of uncertainties and discrepancies among the measurements. The quality of data is 

graded from A to F as follows: 

A 	 Excellent statistical data base providing a clear picture of the 

global distribution of the particular species with uncertainties 

of 5 percent or better 
93 

Compound 

N2 0 

CH3 CI 

CF 2 C12 (FI2) 

CFCI3 (FI1) 

CCI
4 


C2H3C13 


CHCI2F (F21) 


C2 C13 F3 (F113) 


CHC13 


CH 3 1 


CH 3 Br 

C2 C14 

C2 HCI3 

SF 6 

CH4 


CO 2 

C2 H6 

H20 

CO 

H2 


A anthropogenic 

N = natural 

MINOR SPECIES AND AEROSOLS 

Table 26
 

Tropospheric Sources
 

Source 


N 

N 

A 

A 

A 

A 

A 

A 

A 

N 

N 

A 

A 

A 

N+A 

N+A 

N+A 

N 

N+A 

N 

ppbv 

pptv 

ppmv 

= 

= 

= 

Quality 

B 

C 

B 

B 

C 

C 

C 

C 

C 

C 

C 

C 

C 

C 

B 

A 

B 

A 

B 

A 

Mixing Ratio 

290-330 ppbv 

500-750 pptv 

200-230 pptv 

110-140 pptv 

100-150 pptv 

" 100 pptv 

< 10 pptv 

'- 20 pptv 

- 20 pptv 

" 10 pptv 

< 10 pptv 

- 30 pptv 

"-20 pptv 

- 0.2 pptv 

- 1.6 ppmv 

320-331 ppmv 

< 50 ppbv 

Highly variable 

< 100-400 ppbv 

0.4-1.3 ppmv 

parts per billion per volume 

parts per trillion per volume 

parts per million per volume 
94 CHLOROFLUOROMETHANE WORKSHOP REPORT 

B 	 Fair data base yielding statistically useful results, but with uncertainties equal to or exceeding 15 percent 

C 	 Fragmentary information with incomplete coverage 

D 	 Single observations 

F 	 No measurements 

Recent developments ofimportance are: 

Continuing research by groups such as those of Schmeltekopf at the 

Aeronomy Laboratory, National Oceanic and Atmospheric Administration (NOAA); Rassmussen at Washington State University 

(WSU); Inn at the Ames Research Center; National Aeronautics and 

Space Administration (NASA); and Singh at Stanford Research 

Institute (SRI) has solidified the tropospheric data base significantly. 

The lifetime of CH 3 CC13 in the troposphere determined by measured mixing ratio and known source strengtis is a factor of 5 

greater than the lifetime inferred from 1-D model OH concentrations and the rate constant for OH + CH3 CC13 (Singh, 1977). This 

suggests that the average tropospheric level of Of is 3 X 10Scm-3 

to 5 X 10 cm 0 which is significantly lower than that calculated 

from existing models (Singh, 1977). Measured values of OH range 

from 7 X 106 cm- to less than 2 X 106 cm 3 (Wang et al., 1975; 

Perner et al., 1976; and Davis et al., 1976). When the results of 

Davis et al. (1976) are scaled for seasonal and diurnal effects, a 

cm -3 value of 9 X 105 is obtained for the globally averaged OH 

concentration (National Academy of Sciences, Appendix C, 1976). 

When consideration of possible uncertainties in the OH measurements is made, it is unclear whether a real difference exists between 

the reported global values of OH concentration. However, the 

lower value for the global OH concentration might be significant 

because current assumptions regarding the troposphere's ability to 

eliminate moderately reactive chlorofluorocarbons, and theories regarding CO2, N2 0, and methane would require reexamination. 

Uncertainties in the average mixing ratio of CCI4 , CH3Cl, and 

CH 3 CC13 are unacceptably high, making an accurate appraisal of 

the total chlorine budget in the stratosphere difficult or impossible. 
MINOR SPECIES AND AEROSOLS 95 

A convergence of N.0 data on a mixing ratio of 270 to 350 parts 

per billion per volume (ppbv) has significantly narrowed previous 

uncertainties. However, details of the global distribution of N0 

are crucial for determining the tropospheric lifetimes of this molecule, and, although adequate data are not available, they are of 

major importance. The measurement of source terms in the stratosphere is progressing, but the data are sparse. Table 27 represents 

the inventory of source terms that have been observed. Several 

facts have begun to emerge: 

- The latitude studies of the vertical distribution of N.0, CFC13, 

and CFI2 Cl by groups at NOAA and NASA using aircraft have 

begun to define major differences in the upward transport of the 

stratosphere as a function of latitude. This emphasizes the 

difference between: (1) small-scale "diffusion" driven only by a 

density gradient, and (2) a velocity field that represents macroscopic transport independent of density gradients. It is also clear 

that an apparent upward diffusion coefficient cannot be applied 

to the apparent downward diffusion of, for example, a sink 

term. 

- It is desirable to extend the measurements a significant distance 

above the tropopause in order to clearly establish the mixingratio gradient as a function of altitude. 

- Accuracy is limited by the absence of careful cross-calibration 

of gas chromatograph detectors. 

- More data on methane through the stratopause are needed if 

improved vertical eddy-diffusion coefficients are to be inferred 

from such information. 

SINK TERMS 

The major stratospheric photochemical sink terms are HC], HNO3, H0, 

N2 0, CIONO2 , COCd2 , COFCI, and HF. Note that, although some of 

these species can be regarded as a source term under certain conditions 

(e.g., H 0 and HC1), in a time-integrated sense, continuity requires the 

net downward transport of these species from the stratosphere. By 

definition, all sink terms are the product of radical recombination steps 

and are at the same time the reservoir from which the atoms and radicals 

are formed by photolysis and/or oxidation. 
96 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Table 27 

Stratospheric Sources 

Compound Quality Altitude Range 

N20 B Troposphere-43 km 

CH3 Cl C Troposphere-19 km 

CF2a2 (Ff2) B Troposphere-43 km 

CFC 3A'( FI) B Troposphere-43 km 

CCI4 C Troposphere-19 km 

CH 4 C Troposphere-43 km 

CO2 C Scattered 

H20 B Troposphere-32 km 

CO C Scattered 

"H2 C Troposphere-43 km 

Figure 10 summarizes the available HCl data that result from infrared observations and filter experiments. In brief, the measurements show the 

HC1 mixing ratio (by number) to increase from approximately zero parts 

per billion (ppb) at the tropopause to 1.0 to 1.5 ± 0.5 ppb at 28 to 30 

kilometers. ,Uncertainties in the infrared (IR) measurements increase from 

± 25 percent in the lower stratosphere to ± 50 percent at 28 kilometers, 

resulting primarily from diminished absorption. Uncertainty remains regarding the absence of the R5 line of H35C1 absorption spectrum. The 

average of four seasonal determinations of HC1 by the impregnated-filter 

method yields results in reasonable agreement with results from spectroscopic techniques, although it is generally slightly lower. 
97 

50 

0 

MINOR SPECIES AND AEROSOLS 

40I 

lA 

Li 

uJJ 3v 

S20 

0 S.C'SCALCUtATION 

A ACKERMAN ETAL (197&) 

10 M WiLLIAmSETAL. (I9D6 

D F. FARMER ETAL (19"6} 

L LAZRUS ETAL (1975 

0 0510 Is2 

MOLE FIACTION (10.9 

Figure 10. Measured hydrochloric-acid mixing ratios. 

Data of Lazrus etal. (1975 and 1976) above 25 kilometers show a decrease in 

HCl to 0.49 ppbv at 31 kilometers that increases to 0.99 ppbv at 36 kilometers. An assessment of the experimental accuracy (as distinct from 

precision) of these results is currently under study and involves establishing the sensitivity of the technique to species such as CIO, CIONO 2 , 

and COFC1. If such species are contributing to the observed acidic 

chlorine, the effect would be to decrease the apparent HC1 densities. Filter collection efficiencies for acidic C1 have been tested by tandem filters, 

and no acid chlorine is observed on the second filter. In addition to 

midlatitude vertical profiles of HCI, its latitude distributions are being 

studied by using the impregnated-filter method, aboard aircraft. From 

these data, stratospheric inventories are obtained that are used to compute 

the fluxes of HC out of the stratosphere. Initial measurements at different latitudes indicate that concentrations of HCl are lower in equatorial regions, at least during spring. 
98 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Nitric-acid profile and column measurements have been made by filterpaper absorption (Lazlus and Gandrud, 1974a), by IR and submiillimeter 

emission (Murcray et aL, 1975; Harries et al., 1974; and Evans et a., 

1976c); and by infrared absorption (Fontanella et al., 1975). Stratospheric altitude profile measurements usually display a more sharply 

peaked or multipeaked distribution in the lower stratosphere than that of 

model simulations. Various profile and total-column measurements made 

over a few days often exhibit differences on the order of a factor of 2, 

which 1-D models cannot simulate. On the other hand, data obtained 

from balloon platforms in summer at 510 to 590N for 3 successive years 

demonstrate remarkable consistency with only slight changes in the altitude of the peak mixing ratio (Evans et al., 1976c). Although profile 

measurements show no unequivocal seasonal trends, the existing data may 

suggest larger amounts of HNO3 during winter/spring than during summer/fall. Winter measurements at northern and southern latitudes are 

needed to substantiate this possible trend. 

Stratospheric column measurements of HNO3 (Mureray et a., 1975) indicate strong increases as polar latitudes are approached in both hemispheres during January and April Profile measurements (Lazius and 

Gandrud, 1974a; and Murcray et al, 1975) in early spring and late fall 

are usually not inconsistent with these observations. In contrast, column 

measurements by Harries et al. (1974) exhibit no obvious dependence on 

latitude or season. The reason for these differences is not known. However, the latter measurements were generally made over a small latitude 

range at widely spaced times so that short-term fluctuations may have 

masked the trends observed by Murcray et al. (1975). One flight of 

Harries et al. (1974) in June between 500S and 500 N indicates no latitude 

dependence. Perhaps the latitude dependence of HNO3 is strong only 

during January to April. Further measurements are needed to clarify this 

situation. 

Nitric acid is not expected to have a very significant diurnal variation 

(Evans et al., 1976a). 

For comparison of observations with 1-D model simulations, most of the 

available profile data have been grouped between latitudes approximately 

300 to 450N and approximately 510 to 64 0N and have been averaged to 

give figure 11. All of the data were determined between mid-April and 

mid-November. Because the altitude of the peak mixing ratio varies, the 
MINOR SPECIES AND AEROSOLS 99 

SQ I I I 1I1i11i I I I 1 1I 

34

3? 

S I
 

24
 

o25 

i14
 

MIXING RATIO fppbv) 

Figure 11. Calculated average of stratospheric measurements of
 

nitric-acid mixing ratios (Ridley and Wuebbles, 1977).
 
100 CHLOROFLUOROMETHANE WORKSHOP REPORT 

averaging procedure smears the rather sharp layer structure observed in 

individual profiles. It is also evident that the average results show no 

strong latitude dependence despite the two rather wide latitude groupings. In contrast, profile measurements near the Equator are generally 

significantly lower than higher latitude measurements (Lazrus and Gandrud, 1974a). It would be more appropriate to form these averages 

referenced to the height of the tropopause. This was not done, however, because the tropopause information is not always reported for 

the NO measurements. 

King et al. (1976) were unable to identify absorption features attributable to N2 05 in the twilight absorption spectra of Murcray et al. (1969), 

but assigned upper limits of approximately 0.5 ppbv at 30 kilometers, 

decreasing to approximately 0.1 ppbv at 20 kilometers. On the basis of 

radiometric observations in 1974, Evans (private communication, 1977) 

assigned an upper limit of 0.5 ppbv below 30 kilometers. Both observations were attempted (for other reasons) late in the afternoon when N2 05 

is expected to be at a minimum in the diurnal cycle. More recently, 

Evans (private communication, 1977) has assigned an upper limit of 0.5 

ppbv below 30 kilometers from a balloon ascent just after noon. This 

would indicate a high quantum yield for N20, photolysis. 

If the difference between the NO2 sunset and sunrise profiles of Kerr and 

McElroy (1976) is attributed to the change in N2 0 abundance at night, 

then AN2 0 5 = ANO2 /2 - 1 ppbv in the 24- to 32-kilometer region at 

510 N in summer. Combined with the upper limit measurements, the 

maximum expected N205 mixing ratio would be approximately 1.5 

ppbv. 

Chlorine nitrate remains at the center of controversy. Infrared spectroscopy appears to be the most promising method for establishing the 

mixing ratio of CIONO2 between 20 and 30 kilometers, "4hich is the 

expected altitude interval of its peak mixing ratio. Upper limits have been 

placed on the amount of CIONO 2 present in the stratosphere by Murcray 

et al. (1977). Without an accurate knowledge of the ClONO 2 mixing ratio 

between 25 and 30 kilometers, it is difficult to interpret the observed HC1 

concentrations in terms of total chlorine mixing ratios. 
MINOR SPECIES AND AEROSOLS 101 

Extensive work on H20 has been recently reviewed by Harries (1977), 

who offered the following conclusions: 

* 	 Measured water-vapor mass mixing ratios are generally found to lie 

in the range 0.5 to 5 mg kg"1 for all levels in the stratosphere up to 

about 40, kilometers. Higher values have been reported on a few 

occasions that, if they are corroborated, may be an indication of 

temporal variations. Some evidence exists that the mixing ratio 

increases slowly with height between the tropopause and this level. 

" 	Within the quoted range, however, there exists considerable divergence among, the results of different workers for apparently similar 

locations and times of observations. At least part of these differences would appear to be instrumental in origin. 

* 	It therefore seems imperative that Workers, either by personal initiative or under an international program, should combine'to compare 

their instrumental sensitivities and errors so that some standardization of results may be achieved. Without such standardization, 

much information about the real nature and behavior of stratospheric humidity is being lost. The purpose of an experiment 

should no longer be to decide whether 'or-not the stratosphere is 

"dry," but should be to identify small variations of an already 

small relative humidity in order to study the detailed life cycle of 

the vapor. 

* 	By adopting a pragmatic approach, current summaries have been 

produced of the variations of stratospheric humidity with height, 

season, latitude, and time from the data available in the literature. 

The results' are as follows: 

- In midnorthern latitudes, the vertical profile of humidity shows 

a fairly constant mixing ratio between 13 and 30 kilometers of 

about 2.5 ± 0,7 mg kg" . Some evidence for a small increase 

above 30 kilometers exists, to about 2.9 Ing kg-t on the average, 

but perhaps as high as 4 or 5 mg kg"' on occasion. 
102 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Certain results indicate very low mixing ratios in north polar air, 

but these require confirmation. Other data indicate, on the 

average, lower mixing ratios in the lower stratosphere at the 

southern winter pole than at the northern summer pole. 

Similarly, limited evidence of weak, shallow layers of water vapor 

at about 25 to 30 kilometers (within which the mixing ratio rises - to approximately 4 mg kg ' over 1 or 2 kilometers) exists but 

requires independent confirmation. These layers are much too 

high to be associated with the midlatitude tropopause breaks. 

The latitudinal distributions of lower stratospheric water vapor 

obtained by several workers show a general trend to a maximum 

near the Equator, the mixing ratio decreasing at higher latitudes 

in both hemispheres. There is considerable uncertainty, however, 

as to the absolute levels of humidity and also as to trends at latitudes greater than 30°N and 300S. Data are particularly sparse in 

the Southern Hemisphere. 

The temporal variations of water vapor at various levels above 

single locations show marked annual cycles: the amplitudes of 

such cycles obtained by different workers seem to be in fair quantitative agreement. Such cycles may be related to annual variations in the tropical cold-trap temperature. In addition, three independent sets of observations show linear increases in the mean 

- " level of the mixing ratio between 0.1 and 0.25 mg kg ' yr , 

which may possibly be related to long-term changes of temperature at the tropical tropopause. 

The use of experimentally measured values of stratospheric humidity 

in atmospheric models has been investigated. It has been shown 

that extreme care must be exercised in the choice of values to use and 

that such considerable uncertainties exist that quantitative arguments 

that are based on these values may contain considerable uncertainties themselves. This is acceptable only as long as the implications of 

adopting a real experimental result, and the hidden errors that it may 

contain are fully appreciated. 

The variability in OH observed by Burnett (1976) has now made it necessary to determine the in situ OH/H 2 0 ratio in order to determine if fluctuation in H2 0 or if some other species (perhaps 03) is responsible. 
MINOR SPECIES AND AEROSOLS 103 

The HF molecule has a strong infrared absorption feature at 4039.0 enf, 

which is a region that is relatively free from interference by other atmospheric absorption features. Zander (1975) observed this region of solar 

spectra from aballoon platform. This spectrum was obtained at a relatively 

high solar-elevation angle. As a result, it is possible that solar HF may be 

contributing to the absorption. Zander has also observed the absorption 

in spectra obtained from the ground at low solar-elevation angles. These 

spectra are obtained under conditions in which any solar contribution 

would be insignificant. However, the data yield only total-column density. 

By comparing the two spectra, Zander estimates a mixing ratio of 0.3 ppbv 

near 28 kilometers. These results were recently confirmed by Zander et al. 

(1977) from a second balloon flight. However, if solar absorption is occurring in the observed spectra, then the mixing ratios should be considered 

to be an upper limit. 

Although this same spectral region has also been observed during sunset by 

Buijs (private communication, 1976) on several flights made in 1976, 

results have not yet been analyzed. 

Lazrus (private communication, 1976) has also measured HF using impregnated-filter techniques. These measurements yield HF mixing ratios 

on the order 0.1 to 0.2 ppbv. Recent measurements should yield latitude 

and altitude profiles in the Northern Hemisphere for HF. Reduction of 

these data isin progress. 

RADICAL SPECIES 

A major movement is underway, shifting from the problem of verifying 

the existence of the major stratospheric radicals (which are here taken to 

include NO, N02, OH, H02, Cl, and CIO) to the simultaneous observation 

of these species. This change of emphasis results from rapidly expanding 

measurement techniques, observed variability of these short-lifetime species, and a continuing appreciation for the strong chemical coupling 

among the NOx, HOx, and C10, systems. Several other atoms and radicals 

are ofimportance-N0 3, N, H, OC10, and CO2 ; but, given finite resources, 

those species that are believed to control the rate of odd-oxygen destruction must receive priority consideration for the limited task of CFM ozonedepletion assessment. 

Surprisingly few profile measurements of NO2 have been made from 

stratospheric platforms, considering the availability of quite clean absorption lines in the ultraviolet (UV) and IF, No in situ measurements have 
104 CHLOROFLUOROMETHANE WORKSHOP REPORT 

been reported. Instead, profiles have been derived from long-path twilight 

absorption measurements predominantly at sunset (Kerr and McElroy, 

1976; Ackerman and Muller, 1972; Ackerman et al., 1975; Fontanella et 

al., 1974; and Mureray et al., 1974) and from radiometric measurements 

(Chaloner et al., 1975). Column measurements using the Sun, Moon, and 

zenith sky have been made from the ground and from aircraft (Brewer et 

al., 1974; and Noxon, 1975 and 1976). 

Profile measurements of NO2 fall into two fairly distinct groups. Pressure 

modulation (Chaloner et al., 1975) and submillimeter (Harries et al., 1976) 

radiometric methods give substantially larger NO 2 mixing ratios during the 

day than those of UV and IR absorption methods at sunset (Kerr and 

McElroy, 1976; Ackerman and %Muller,1972; Ackerman et a., 1975; Fontanella et al., 1974; Murcray et al., 1974). The disparity is larger at night, 

a few hours before sunrise, when the radiometric methods yield mixing 

ratios in the 20- to 35-ppbv range between 25 and 35 kilometers. Current 

theory suggests even larger values at night after sunset to allow for N205 

formation. Thus, the long-path IR or UV absorption methods are generally 

three to four times lower at comparable times. The daytime NO2 values 

determined radiometrically are larger than would be calculated from NO 

measurements, and the nighttime values are larger than the total NOx predicted by current atmospheric models. The nighttime radiometric profile 

measurements give much larger total column NO2 than those measured 

directly by Noxon (1975 and 1976). Either systematic errors in the different techniques or in data reduction are suspected, and in-flight NO2 instrument intercomparisons are required. 

If the radiometric measurements are neglected for the moment, the remaining published NO2 profiles do not exhibit the order of variability observed for HNO 3 and NO. This may be attributable to the more limited 

data base. Variations in total-column NO2 by a factor of 2 or so during a 

month are not uncommon (Noxon, 1975 and 1976). However, changes in 

the lower stratosphere strongly weight column measurements; a 1-kilometer 

layer of 0.2 ppbv at 15 kilometers is equivalent to one at 30 kilometers 

containing about 2 ppbv. 

Profile measurements have been insufficient to determine if seasonal and 

latitudinal trends persist at all altitudes. However, stratospheric column 

measurements by Noxon (1976) indicate a summer-to-winter ratio of about 

2 south of 500N, and about 6 north of 550N. During the summer, little 

change in total NO2 occurred between 350 and 80 0N. During winter 

months, in contrast, the latitude dependence is very strong with about a 

factor of 6decrease in total NO2 above about 500N. Similarly, unpublished 
MINOR SPECIES AND AEROSOLS 105 

experiments by McElroy and Kerr (private communication, 1976) indicate 

significantlylower total NO2 in Arctic air masses that have reached southern 

Ontario. It has been argued qualitatively that this winter decrease is attributable to conversion of NO2 into N2 05 and/or HNO3 during the long 

winter nights. However, a 1-D diurnal model has been used to attempt a 

simulation, but the calculated column decrease was not nearly as great as 

that observed even when the unknown quantum yield of N2 0, photolysis 

was decreased to a reasonably low value (McConnell and Ridley, unpublished study). By default, then, the tentative conclusion was reached that 

macrotransport processes are responsible. 

Altitude profile measurements of NO have been made by long-path IR absorption at sunset (Ackerman et al., 1975 and 1973), by long-path emission (Chaloner et al., 1975), and by in situ chemiluminescence instruments 

(Ridley et al., 1974, 1975, 1976 and 1977; and Ridley and Bruin, 1977). 

Particular altitudes have been sampled by a laser absorption method (Patel 

et al., 1974; and Burkhardt et at, 1975) and also by chemiluninescence 

methods (Loewenstein and Savage, 1975; and Loewenstein et al., 1975). 

As is the case for HNO3 data, in situ and long-path techniques exhibit a 

spread between individual measurements of up to an order of magnitude in 

the 17- to 30-kilometer region between about 30' n d 450N. Short-term 

fluctuations of several fold are observed in aircraft measurements. In contrast, five balloon profiles in summer between 51 and 590N are quite consistent, as were NO2, HNO3 , and O profiles (Evans et al., 1976a; and 

Ridley and Bruin, 1977). 

Recently, NO chemiluminescence methods have extended measurements 

to 60 kilometers. In June, Drummond et al. (1976) obtained 10.5-ppbv 

constant between 40 and 45 kilometers. In April, Mason and Horvath 

(1976) obtained upper imits roughly two to three times lower in the 

same altitude range. This would be consistent with the observed seasonal 

changes in the 300 to 450N region at lower altitudes, although more measurements are required to determine if the differences are not due to 

natural variability. 

Aircraft measurements (Loewenstein and Savage, 1975; and Loewenstein 

et al., 1975) near 18 and 21 kilometers indicate typical summer values two 

to three times greater than in winter in the 30 to 450N range. Whether 

this trend extends to higher altitudes is not determinable from the balloon 

measurements. 
106 CHLOROFLUOROMETHANE WORKSHOP REPORT 

The latitude trends are also known only in the 18- to 21-kilometer region. 

Between the Equator and approximately 500N, the data of Loewenstein 

and Savage (1975) obtained during summer and fall exhibit only small 

changes. Above approximately 500N, there is a reasonably clear increase 

of up to a factor of about 3. Because profile measurements indicate a substantial increase in mixing ratio with increasing altitude, the latitude observations may be primarily due to deeper penetration of the stratosphere at 

northern latitudes. 

The spread in measurements in the range of 300 to 450 N latitude now appears uncomfortably large. In general, the chemiluminescence measurements (Ridley et al., 1974 and 1975) are lower than those obtained by 

other techniques, although some measurements are in agreement with the 

seasonal trends at lower altitudes. In terms of the present data base, the 

variations appear much greater between 300 to 450 N than between 510 to 

590N, and this may be partially associated with changes in the strong jet 

stream at lower latitudes. Most of the balloon chemiluminescence results 

were obtained well before local noon, and it is now known that NO can 

continue to increase by varying amounts as the solar zenith angle decreases 

from about 900 to the value at local noon (Ridley et al, 1977; and Patel 

et al., 1974). In summary, if systematic errors between various techniques are not discovered, seasonal, diurnal, and transport changes may be 

responsible for the dispersion in the data. Certainly, the major NO, constituent below about 25 kilometers is HNO3 , and separate experiments 

made over a short period of time exhibit large differences. Thus, it is not 

unexpected that NO also exhibits large variability. 

NO3 is expected to have a strong diurnal variation with maximum concentrations during the night. However, nighttime and daytime total-column 

measurements by Noxon (1975) and Davies (1974) failed to identify absorption attributable to NO3 , and only an upper limit of about 5 percent 

of NO2 was set. This is consistent with model simulations; NO3 is rapidly 

photolyzed during the day and slowly converted to N2 0 5 during the night. 

Strong diurnal variations are expected in the concentrations of NO2 , NO, 

NO 3 , and N2s, but this has been confirmed directly by experiment only 

for NO and NO2 Indirect evidence exists for the behavior of N2 0 5 . 

For solar zenith-angle (x) variation between approximately 900 and 950, 

NO2 is rapidly converted to NO during sunrise and NO is converted to 

NO2 during sunset so that NO + NO2 should be essentially constant. At 

sunset and below about 40 kilometers, for example, the decay of NO is 
MINOR SPECIES AND AEROSOLS 107 

governed by the time constant, t = l 1 (JNo + ko+NO2 [0)), where JN0 is 

the photodissociation coefficient for NO2 . JN dominates this expression so that net NO-to-NO2 conversion essentially parallels the attenuation 

of the UV flux. For NO, the rapid disappearance at sunset and appearance 

at sunrise has been measured with a time resolution that is shorter than t in 

the range X = 90 to 950 (Ridley et al., 1976 and 1977; Ridley and Bruin, 

1977; Patel et al., 1974; and Burkhardt et al., 1975). When local ozone 

and temperature data are available, diurnal models can simulate the NO 

behavior quite well (Ridley et al., 1976 and 1977). This type of experimental data provides a stringent test of model chemistry since large-scale 

transport is negligible over the time scale of the experiments. 

The variation of NO has not been measured for a 24-hour period during a 

single balloon flight. However, Burkhardt et al. (1975) have covered the 

period from before UV sunrise until about local noon and from before 

noon until UV sunset with two balloon flights near 28 kilometers at different times of the year. After the initial rapid increase in NO due to NO2 

photolysis at sunrise, the laser absorption measurements exhibit a further 

strong increase in NO until about noon. Similar results were mentioned 

by Chaloner et al. (1975). 

Chemiluninescence measurements (Ridley et al., 1977) during sunrise until 

before noon also showed an increase of NO as Xdecreases from about 90 , 

but the increase was considerably smaller than that of the laser measurement. However, a diurnal model that attributes the morning increase in 

NO to N2 05 photolysis can explain both sets of data if a judicious choice 

of 03 concentration and total NOx is made (Ridley et al., 1977). Measurements of 03 were not made on either flight. For the period of several 

hours after local noon, the NO concentration remains quite constant according to the results of Burkhardt et al. (1975). Chemiluminescence 

measurements on four flights near 35 kilometers have established that, at 

this altitude, the net decay of NO at sunset begins abruptly as x increases 

beyond about 900 (Ridley et al., 1976; and Ridley and Brin, 1977). Only 

a slight decrease in NO is observed as X increases from -.800 to 900 in 

agreement with only a slight decrease in JNo at this altitude. This is in 

apparent disagreement with the afternoon fligt of Burkhardt t al. (1975) 

on which they observed a large reduction in NO beginning near x = 670. 

However, their observations are perturbed by a severe loss in balloon altitude beginning several hours after local noon, and it appears that their 

data are essentially profile measurements intermixed with the diurnal decay 

as sunset approaches. Consequently, there remains a period from midaftemoon until a short time before sunset in which the diurnal behavior of 
10 CHLOROFLUOROMETHANE WORKSHOP REPORT 

NO was not observed. This is unfortunate because these observations of 

NO could contribute to resolving the controversy concerning the possibility of a strong afternoon maximum reported for NO2 (Brewer et al., 

1973 and 1974). 

The diurnal variation of NO2 is less clear because of conflicting reports. 

Long-path solar absorption measurements of NO2 have been made by Kerr 

and McElroy (1976) from a balloon near 35 kilometers during morning = and evening twilight. These measurements are weighted toward x 900 

and show from approximately 0.5 to 2 ppbv more NO2 at sunset than at 

sunrise over the tangent altitude range of 16 to 30 kiometers. Total-column measurements by Noxon (1976) are consistent with this observation. 

These observations alone can be explained by slow conversion of N0 2 to 

N20 via NO3 at night, followed by N2 0 s photolysis to NO2 during the 

day. 

The daytime behavior of NO2 remains in a state of controversy. Groundbased measurements (Breweretal., 1973 and 1974) that have been reduced 

to give vertical proffles indicate a strong afternoon maximum with NO2 

being 3 to 4 times either sunrise or sunset values at all altitudes. In contrast, Noxon (private communication, 1976) cannot confirm this although 

the basis of the experimental method is the same. On the other hand, 

Chaloner et al. (1975) report that daytime NO2 concentrations are lower 

than those after sunrise by up to a factor of 2 at 35 kilometers. Their daytime NO + NO2 values are also much lower than the NO2 measured several 

hours before sunrise. Both of these observations are inconsistent with 

morning measurements of NO and cannot be explained by current stratospheric chemistry. Their nighttime observations would require an additional source of NO2 or NOX. Substantiation of even the relative trends 

observed with this technique is necessary. Clearly, the diurnal behavior 

of NO2 requires further investigation by independent measurement techniques. Present NOx diurnal models below '-40 kilometers show NO 2 decreasing during the night by conversion to NO 3 and N2O0, a rapid decrease 

during sunrise due to photolysis, a slow growth during the day as a result 

ofN2 05 photolysis, and a rapid increase during sunset as NO is oxidized to 

NO2 . The growth during the day depends on the assumed photolysis rate 

of N.0 s . Noxon's (1976) column measurements and NO measurements 

are generally consistent with this scheme. Both daytime profile observations of NO2 cannot be correct. However, if either one is substantiated, 

there must be missing chemistry in the model simulations. 
MINOR SPECIES AND AEROSOLS 109 

Simultaneous measurements of at least 03, temperature, and several NOx 

constituents provide stringent tests for model simulations. Few such measurements have been made. Using their absorption technique at sunset, 

Ackerman et al. (1975) obtained vertical NO and NO2 profiles in May 

near 44°N. This permits a determination of the N0 2!NO ratio that is 

expected to be in photochemical steady state. However, this ratio is quite 

sensitive to local 03 and temperature, and, unless these are measured 

during the flight in the neighborhood of the tangent altitude, a strict 

model comparison is not possible. Chaloner et al. (1975) also determined 

NO and NO, profiles during the day. If possible systematic errors in 

their technique cancel, they observe an NO2 /NO ratio of about 1.6 to 2 in 

the 20- to 33-kilometer region. Again, 03 and temperature data were not 

reported. 

Simultaneous measurements of 03, temperature, NO, NO 2 , and HNO3 

have been made yearly during the summer since 1974 (Evans et al., 1976b 

and 1976c). Profile measurements have been compared with a 1-D model 

that did not include CIO. (Evans et al., 1976a). Although the comparison 

is reasonable, differences of a factor of 2 to 3 exist between the measured 

and calculated NO2/NO, HNO3/NO, and HNO3 NO ratios at some attitudes. Because the latter two ratios depend on the rate coefficient for the 

OH + HO2 reaction, which is not well-known, the differences are not unexpected. More disturbing is the measured NO2/NO ratio, which is 

roughly a factor of 2 greater than that of the model ratio at most altitudes. 

This could be due to several factors. First, either or both measurements 

may have unrecognized systematic errors. Secondly, the model-calculated 

photodissociation coefficients may not reproduce those of the flight, or 

they may have insufficient chemistry. Indeed, if recent CI0 measurements 

( G. Anderson, private communication, 1977) are substantiated, the 

calculated NO2 INO ratio would be significantly increased over that calculated with the measured amounts of C0 neglected Thirdly, the fleas- ­

ured N0 2/NO ratio is a combination of results from an in situ and longpath technique. Thus, the NO measurements were made in an air mass in 

which temperature and 03 were measured, whereas the NO2 measurements 

were made in different air masses that extend farther and farther from the 

balloon environment as the zenith angle increases beyond 90O

° 

Several independent methods are now being applied to the measurement of 

"reactive" chlorine, designated here as CIOX, which is taken to include 

C1, C0, OCIO, and CIO2 . Of those four species, only the first two are 

believed to play a major part in the definition of Ntratospheric chlorine 

photochemistry, and thus they will receive the major amount of attention. 

However, it must be realized that no measurements of OCIO or C1O 2 exist, 
110 CHLOROFLUOROMETHANE WORKSHOP REPORT 

and thus their exclusion from further comment is not based on empirical 

information. 

Chlorine monoxide is now under investigation by four independent 

methods: (1) solar ultraviolet absorption, (2) microwave emission, (3) cryogenic entrapment, and (4) in situ chemical conversion-resonance fluorescence. Atomic chlorine has been studied by using in situ atomic resonance 

fluorescence only. For chlorine monoxide, techniques (1) and (2) have 

yielded upper limts; technique (4) has been used to observe the altitude 

distribution. 

Three simultaneous observations of Cl and CIO have been performed using 

the in situ resonance fluorescence technique, the first on July 28, 1976, at 

12:00 noon Central Daylight Time (CDT) (local solar zenith angle 16), 

the second on October 2,1976, at 12:00 noon CDT (local zenith angle 350), 

and the third on December 8, 1976, (local solar zenith angle 500). In all 

cases, a 1.5 X 107 ft3 helium balloon, launched at dawn from Palestine, 

Texas, 32 0N latitude, bore the experimental package to an altitude of 

43 kilometers (141,000 feet). After activation and stabilization ofinstruments, the measurement phase began on deployment of the "guide-surface" 

parachute, which controlled both the velocity and attitude of the package 

during descent. The concentrations of both Cl and CIO were successfully 

measured on each flight with the results shown in figure 12. Significant 

vertical structure was apparent, particularly in the CIO profile for which 

the signal-to-noise ratio exceeded 10'2 throughout the region between 25 

and 40 kilometers. In the case of atomic chlorine, the signal-to-noise ratio 

exceeded 3 above 40 kilometers, decreasing to 2 at 35 kilometers, below 

which contributions to the background count rate in the absence of Cl 

exceeded the fluorescence signal. It is expected that data extending to 

concentrations less than 105c= 3 for Cl (7P) will be feasible on future 

flights. An experimental uncertainty of ±40 percent is assigned to these 

data based partially on laboratory absolute calibration uncertainties 

(±20 percent) and partially on a subjective appraisal of difficulties inherent in the control of any such field experiment. 

Figure 13 presents a comparison between the data and four theoretical 

predictions of C1 and CIO appropriate to conditions during the flights. 

The difference in solar zenith angle between the July and October data is 

not a significant consideration, although it may be important with respect 

'to the December data. It is impossible to compare these experimental results with all the available theoretical predictions; the foregoing four models 

were selected because: (1) they represent a cross section of the major 
OATA AVERAGED OVER I-NANOMETER INTERVALS 

45 CI I P) CID 

C)C 

30 320N 

C> 

25 

t 

SOLAR ZENITH4 UNqERTAI!Ty 

SYMQOL DATE ANGLE LOCALTIME 

20 0 28 JULY 76 16 12LPMCDT 150y t0 

S "20CT76 35 12ISFMCDT 5Q% 14% 

o 0EC 76 550 1200PMCST 1w% 235% 

105 106 107 106 1 

CONCENTRATION (cm -

3) 

Figure 12. Measured concentrations of atomic chlorine and chlorine oxide. F-i 
-1 -

L 4 

~wMO
 

­

< 

C) 

F- S ULUCICERONE, AND DONAHUE I LCDI 12 3 mb} . 

W0FSY,MCGLR0Y IWMEIi2,?ppb WM M 

CHANtGE WU BBLES ' {CN] (16,upt) CM LCD 

Z 30 UTZE",MeAFEEICMI111PPb1 

5 7 


10 100I look-


CONCENTRATION (CNI2 

Figure 13. Comparison of measured and theoretical concentrations of atomic chlorine and 

ch[orine oxide. 
MINOR SPECIES AND AEROSOLS 	 113 

theoretical efforts to quantitatively predict chlorine catalyzed ozone depletion, and (2) they adequately encompass the range of theoretically predicted Cl and C1O concentrations. These-determinations were made prior 

tothe new determination of Howard (1977) for the reaction of HO2 + NO. 

The theoretical consequences of employing the new rate are discussed in 

Chapter 4. 

Several conclusions can be drawn from figures 12 and 13: 

* 	 Cl (2 p) and CIO exist in the stratosphere. 

" 	 Significant variabilityin the CIO and Cl concentrations exists throughout the stratosphere. 

* 	 Significant local structure is apparent in the C10 densities for which 

the signal-to-noise ratio is approximately 102. 

* 	 The ratio of C10 to CI is approximately 102 at 40 kilometers and 103 

at 35 kilometers. 

* 	Cl and C10 densities are equal to or greater than those predicted by 

theory. (A full discussion of these results appears in Chapter 4.) 

Although several oxygen-hydrogen radicals exist, OH is the only species 

for which measurements 'have been reported in the troposphere and the 

stratosphere. OH is generally considered to be a key hydrogen radical 

linking the NO, ,HOx, and ClO, families. Since the NAS report, groundbased, column-density measurements of Burnett (1976) have become 

available. 

Burnett's (1976) results may be summarized by noting that he observed 

column-OH concentrations as high as 4.4 X 1014 cm "2 and fluctuations 

in the column greater than factors of 3 in a few.hours. Figure 14 summarizes all existing OH measurements to date. The tropospheric measurements 

of Wang et al. (1975) and Perner (1976) may not be representative of the 

natural troposphere because they are made at ground level. Taking the 

maximum value inferred from the Davis (1976) measurements, the column 

content contribution from the troposphere is less than 2 X 1013 cm "2 . 

This means that Burnett's (1976) integrated column content of 4.4 X 1014 

must arise principally from the stratosphere and mesosphere. J. G. Anderson (1971 and 1976) has recently revised his rocket data using new values 
114 CHLOROFLUOROMETHANE WORKSHOP REPORT 

70 

50 

50 	­ STRATOPAUSE


40
 

,
 

0 

I-F 

20 	 ­

.TFOPOPAUSE 

10 

1 I 7 I I 	 1 

106 	 107 108
 

" OH CONCENTRATION (MOLECULES cm 3

) 

NOTE DATA SOURCES ARE AS FOLLOWS 

O 	 ANDERSON (1971), BY MONITORING SCArTERED 306 4-NANOMETER SUNLIGHT OVER WHITE
 

SANDS. NEW MEXICO, IN APRIL WITH THE SOLAR ZENITH ANGLE AT 860
 

E 	WANG ETAL (19751. BY DETECTING LASER-INDUCED FLUORESCENCE CENTERED AT 309 NANOMETERS 

AT THE GROUND NEAR DEARBORN, MICHIGAN, DURING 4 DAYS IN AUGUST 

O 	 ANDERSON (1976). BY MEASURING RESONANCE FLUORESCENCE AT ABOUT 309 

NANOMETERS OVER PALESTINE, TEXAS, IN JULY (S) AND JANUARY AT A ZENITH ANGLE OF 80" 

\ 	 BURNETT (1976). BY COLUMN-INTEGRATED RESONANCE ABSORPTION OF 308-NANOMETER 

SUNLIGHT FROM THE GROUND AT FRITZ PEAK OBSERVATORY IN SEPTEMBER AND DECEMBER 

(THE INTEGRATED OH COLUMN HAS BEEN UNIFORMLY DISTRIBUTED BETWEEN 10 AND 70 

KILOMETERS) 

E 	DAVISE T AL (1976). BY OBSERVING LASER-INDUCED FLUORESCENCE AT 309 5 NANOMETERS
 

AT 210 AND 31' NORTH LATITUDES IN OCTOBER IN DAYLIGHT
 

A 	 IN THE AIR PERNER ET AL (1976), BY LASBER IGHT ABSORPTION AT 307.995 NANOMETERS 

OVER JULICH, GERMANY (SIG NORTH LATITUDE), IN LATE SUMMER AND FALL 

Figure 14. Observed atmospheric hydroxyl radical concentrations. 
MINOR SPECIES AND AEROSOLS 115 

for the rate of quenching of A2E+ with N2 and 02 (Becker, 1972). Applying this revision, plus a factor to account for the time of day, satisfactory agreement can be obtained between Burnett (1976) and J. G. Andersons (1971 and 1976) values. 

Several conclusions follow: 

" 	Column densities of the hydroxyl radical are observed to vary by 

factors of 3 from day to day. 

" 	 Column and in situ measurements are in reasonable agreement. 

* 	 The determination of other radical concentrations that are known 

to depend on OH should be made simultaneously with that of OH. 

Also, OH itself should be measured simultaneously with O and H2 0 

in order to determine the nature of the observed variability. 

SOLAR FLUX AND 00 RATIO, 

Three different types of information related to solar-flux measurements 

are required for a detailed knowledge of the photochemistry of the stratosphere: (1) the solar flux that reaches the top of the terrestrial atmosphere, (2) the solar flux at specific altitudes, and (3) the integrated effect 

in terms of the concentrations of photochenical species with short lifetimes. 

The ultraviolet solar flux in the region from 160 to 300 nanometers is 

responsible for the photochemistry and heating of the stratosphere, which, 

in turn, is an important driving function for stratospheric transport processes. This wavelength region is noted for its complexity in terms of both 

terrestrial atmospheric and solar processes. During the early years of 

measurements of ultraviolet solar flux, it was not uncommon for measurements by independent groups to differ by factors of 2 to 3. Since that 

time, the techniques for measurement and absolute spectroradiometric 

calibrations have improved enormously. 

Physical processes on the Sun that give rise to this radiation are extremely 

complex because the solar spectrum changes from a continuum with superimposed absorption lines at long wavelengths to a continuum with superimposed emission lines at the shorter wavelengths. Correspondingly, observations of the solar disk change their character from limb darkening to 

limb brightening at the shorter wavelengths. 
116 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Of great importance to understanding the physics and chemistry of the 

stratosphere and climatic trends is determining what the values of the 

ultraviolet solar spectral irradiance are and whether it has varied with time. 

A current tabulation of some major sources of data on solar spectral itradiance may be found in Heath and Thekaehara (1976). Figures 15 and 

16 graphically show these data sets. Not included, in these figures are the 

data by Detwiler et al. (1961), which differ significantly from all later 

measurements. Figures 15 and 16 include data from Ackerman (1971), 

Simon (1975), and Heath (1973). 

The other important aspect of the ultraviolet solar flux is the nature of its 

temporal variability. The effects of solar-flux variability for wavelengths 

less than 120 nanometers have been established from the analysis of satellite 

drag measurements for both the 27-day solar rotational period and the 

11-year sunspot cycle by King-Hele and Quinn (1965) and Jacehia (1963), 

respectively. Although the 27-day variability in the ultraviolet solar flux 

below 120 nanometers has been observed directly, the variability for the 

flux for the 11-year period has not. 

Observations of the 27-day variability of the solar flux in the region of 160 

to 300 nanometers have been reported by Heath (1974) and Hinteregger 

(private communication, 1975). The maximum variability for this wavelength region appears to be less than 10 percent per solar rotational period. 

Although it is generally accepted that the solar flux varies over the 27-day 

solar rotational period below 200 nanometers, this is not the case concerning questions of variability of this wavelength region over the 11-year sunspot cycle. Evidence for an 11 -year variability of the Sun from 160 to 300 

nanometers has been given by Heath and Thekaekara (1976). However, 

questions as to the reality of the effect have been raised. Evidence for the 

existence of a solar-cycle effect is derived from measurements on a rocket 

in 1966, on Nimbus-3 in 1969,.on Nimbus-4 in 1970, and on Atmosphere 

Explorer-E (AE-E) in 1975. The ratio of solar flux measured at solar 

minimum to that at solar maximum (shown in figure 17) is approximately 

0.5 at 177 nanometers and approximately 0.8 at- wavelengths near 

290 nanometers. The solar-flux values were derived from both broadband 

photometric observations and from measurements with a double monochromator with a 1.0-nanometer spectral bandpass. There were no measurements at 210 nanometers over the solar cycle. The validity of this 

curve needs to be established by other investigators, particularly in the 

atomic oxygen production region in the stratosphere at approximately 210 nanometers. 
400
 

20 

2 

to 

10I 

o0 00 0 

o 00 

4,I DONNELLY AND POPE 19731 Ca 

0 1, 

0. 

S 0 

' - . ',1 

004 o I ' +

"" i*-

.1 

0 

ORUECI<NER ein. 1ID71) 

ACI(EflM AN 19711 

n,rTMAN 11074) 

ti 

0r 

0

Go 

0 02 

o o0i05 

.o r V SIMON11075) 

.. _ T HeKAEKA RA 073, 1974) 

III 1 I I I 

1400 I00 1000 1700 IBM 009 20 2100 2200 

WAVELENGTH(A. 

Figure 15. Solar spectral irradiance in the wavelength interval 1300 to 2300 A. 

2300 
lOQQ I 

Q 

00 0 

200 

0 

00 

0 6 EROAIPFODT t1972? 

V SIMON 110975 

C 0 THEKAfKARA (113,tSV4J 

o HEATH 173 

O ACKEnMAN 119711 -

I 1 P I I I I I 

2000 2200 2200 2300 2400 2500 26M 2700 00 ?m000 

WAVELENGTH (A) 

Figure 16. Solar spectral irradiance in the wavelength interval 2000 to 3000 A. 
060 

E A 

04 U0 

0 • 197511970 l 

A 1966/1970M 

02 0 

M 1966/1969M 

160 180 200 220 240 260 280 300 320 340 

( nml 

Figure 17. Ratio of solar flux measured near solar-cycle minimum to that measured 

near solar-cycle maximum versus wavelength. 
120 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Unfortunately, not enough is known about the global scale structure of 

the stratosphere over the sunspot cycle to permit the assessment of 

long-term solar variability on the stratosphere from-atmospheric observations. 

In summary, the current measurements of solar spectral irradiance from 

160 to 300 nanometers appear to be in reasonably good agreement for the 

period of solar minimum. A concerted effort should be made to greatly 

increase the number of spectral irradiance measurements as the Sun goes 

from sunspot minimum to sunspot maximum. Emphasis should be-placed 

on improving the accuracy of the absolute spectroradiometric calibration 

of the instruments that are used because it is extremely important to establish whether or not long-term changes are taking place in the ultraviolet solar flux. 

Experimental measurements of the solar flux at specific altitudes in the 

stratosphere are required for verifying the calculated attenuated solar flux 

at that altitude. This is particularly desirable in the Schumann-Runge 

bands. Evens (private communication, 1977) has begun to obtain direct 

integrated solar-flux measurements in the spectral region 180 to 200 nanometers. Megill et al. (private communication, 1977) have obtained solarflux data as a function of wavelength and angle with respect to the line of 

sight with the-Sun. 

The measurement ofphotochemically generated short-lived species provides 

information on the integrated effect of solar flux, photodissociation cross 

sections, and the column densities of major light absorbing species. These 

measurements are important as both altitude-concentration profiles and 

ground-based column densities. Atomic oxygen is a good candidate for 

altitude-concentration profiles. The excited molecular-oxygen species, 

0 2 ("Ag) and 2( Z), are desirable for both column densities and altitudeconcentration profiles. 

The photochemical decomposition of ozone in the 50- to 30-kilometer 

region can be described quite accurately compared to other species. Scattering effects and albedo are calculated to have little influence on this 

process. The-three-body recombination rate constant for forming ozone is 

also one of the better-known kinetic rate constants. The problem in accurately calculating the ozone profile in this region involves the rapidly 

changing lifetimes (2 hours to 10 days) and the destruction by minor species (NOI ClOx, and HOx). For these reasons, simultaneous experimental 

altitude-profile measurements of the local ozone are essential for the determinations of solar flux (X > 200 nanometers), 0 atoms, 0( 1 Ag) , and 02(04) 
MINOR SPECIES AND AEROSOLS 121 

TRANSPORT MEASUREMENTS 

Transport plays an important role in the stratosphere through the mixing 

processes and the circulation patterns. Transport processes are too often 

included only during the hand-waving stage of explaining differences 

between measurements and theory. Knowledge of these processes is essential to the development of useful transport parameters invoked in models. 

Understanding of the stratosphere has reached the point at which transport 

measurements should be made simultaneously with measurements of 

stratospheric composition. 

This subsection describes the types of transport processes that affect composition measurements, the types of transport measurements that can be 

performed, and their accuracy. 

Altitude-profile measurements of stratospheric constituents are performed 

on a mesoscale basis (100 kilometers to 100 meters). At this scale, large 

variability exists due to localized transport processes. The danger exists 

that the constituent measurements are accepted as representative of a global 

scale, which in fact they are not. Furthermore, no attempt is currently 

being made to understand or parameterize the mesoscale variations. Until 

these mesoscale variations are understood, it will be impossible to derive 

representative averages of constituent densities and eddy-diffusion coefficients for l-D and 2-D models. It is established that tropospheric folding 

is a dominant mechanism for troposphere/stratosphere interchange at midlatitudes. Midlatitude measurements of constituents may vary significantly 

if a tropopause folding occurs at the measurement location and times. 

Similarly, in the tropics where Hadley cell circulation dominates the lower 

stratosphere, altitude profiles alone should not be used to infer an eddydiffusion coefficient. It is necessary to separate the effects of convective 

transport on a constituent .profile before attempting to estimate the diffusive transport. This requires knowledge of the velocity field or, in some 

instances, the flux of an effective tracer (such as ozone). Such information 

is not readily available. Horizontal variability on the mesoscale is a result 

of irregular mixing and wind shear. At one altitude, the composition maybe transported from the southwest, whereas, at another altitude, it is from 

the northwest. Hence, a single altitude-profile measurement can represent 

chemical processes occurring at widely separated locations. Thus, it is 

necessary to perform a meteorological analysis to trace the origin of the 

constituent as a function of altitude. 
122 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Frequently, the midlatitude tropopause is not sharply defined and consists 

of multiple temperature minimums with associated sharp increases in the 

ozone mixing ratio. Measurements of ozone taken simultaneously with 

other measurements are extremely useful in defining the tropopause structure and the transport processes above the tropopause. 

The location of measurements is important, and a particular local feature 

can affect the data (e.g., measurements conducted near mountains). High 

Altitude Clear-Air Turbulence (HICAT) program data clearly indicate enhanced turbulence effects near mountains, probably because of breaking 

waves (see Climatic Impact Assessment Program (CIAP) Monograph I, 

Chapter 6). Certain latitudes pose problems attributable to the confluence 

of different circulation patterns. 

Measurements can be conducted at specific locations to minimize or maximize one or the other circulation patterns. Measurements conducted in 

New Mexico and Texas are in the transition region between tropical effects 

and midlatitude effects (i.e., at the tropopause gap). A more effective 

procedure would be to conduct measurements farther south, such as in 

Panama, to isolate tropical transport effects and farther north to examine 

midlatitude transport. 

Frequently, logistics considerations demand that balloon flights be conducted duringthe spring or fall turnaround. Unfortunately, these are times 

of rapidly changing circulation conditions, and measurement results are 

difficult to interpret in terms of being representative of a particular season. 

WINDS AND DYNAMICS 

The measurement of vector winds and dynamical processes in the stratosphere is significantly more difficult than measurement of scalar quantities, 

such as temperature, composition, or density. Because this difficulty has 

limited the number of measurements, present knowledge of the winds and 

dynamics is poor in quantitative accuracy in spite of the fact that transport 

plays a dominant role in the variation of trace constituent densities about 

the globe and the small-scale structure observed in ozone-density profiles 

(Dfttsch, 1969). The general features of transport are known on a large 

scale such as geostrophic flow, two-cell meridional flow, and quasi-biennial 

oscillations. However, detailed measurements of the winds on the synoptic 

and mesoscales are not accurately known over the surface of the globe. 

Assuming that there is a need for detecting a perturbation in the stratospheric environment within a short period of a few years, it will be necessary to improve the present measurements of winds and dynamics. 
MINOR SPECIES AND AEROSOLS 123 

The present measurement techniques can be grouped as either direct measurements or indirect measurements, the latter using numerical models to 

infer a transport term from known distributions of tracers. Most of the 

direct measurements of winds have been performed with rawinsondes and 

rocketsondes, small payloads that are tracked in ascent or descent to 

deduce a wind velocity from displacements of their horizontal position 

with time. These sondes are launched on a regular schedule as part of the 

Cooperative Meteorological Rocket Network (CMRN). Additional highaltitude data have been obtained from observations of improved radiosondes (Gimspheres)and grenade experiments. Winds are also deduced from 

changes in position of balloons, ribbon chaff, falling spheres, and smoke 

trails. Direct measurement of circulation patterns are observed in the patterns of atmospheric tracers such as radioactive debris and ozone. Descriptions of these measurements and typical results canbe found in the excellent summaries in the CIAP Monographs I and Ill. Direct measurement of 

small-scale turbulence has recently been conducted with anemometers 

suspended below balloons (Cadet, 1975; and Barat, 1975). Projects HICAT 

and Coldscan (see CLAP Monograph I, Chapter 6), conducted between 1964 

and 1970, used aircraft to measure clear-air turbulence (CAT). These data 

are significant in describing the probabilities of occurrence, scale, and intensity of turbulence and in correlating the CAT with synoptic structure. 

The primary source of wind measurements in the stratosphere is the 

rocketsonde data obtained at 30 active stations and archived at World Data 

Center-A (WDC-A), Asheville, North Carolina. The data contain systematic 

bias due to the time of observation and systematic errors of each particular sonde and technique of data analysis used, in addition to random instrumental errors. Although WDC-A exercises care to minimize obvious 

data errors, the user usually accepts the mean values as representative, 

which may not be the case. The rocket launchings are usually conducted 

at local noon with the exception of the measurements by the Union of 

Soviet Socialist Republic (USSR), which are conducted at night. The 

predominance of noon launchings places a bias in the mean values originating from the diurnal tide. This is particularly noticeable in meridional 

winds (Nastrom et al., 1975). The infrequency of the soundings and the 

incomplete global coverage create many statistical and analytical problems. 

This is illustrated by comparison of the monthly mean winds whose variance is equal in magnitude to the mean wind speed. This is true even for 

the strong zonal winds, which are observed to have-significant longitudinal 

variations in the upper stratosphere from the monthly mean, such as 

between West Geirnish and Ft. Greely. It is obvious that present wind 

measurements are inadequate to accurately describe the dynamics of the 
124 CHLOROFLJOROMETHANE WORKSHOP REPORT 

stratosphere. The accuracy of the measurements is difficult to determine 

from the limited number of simultaneous soundings. The problem of systematic errom in technique is illustrated by the results of comparative 

rocketsonde tests performed jointly by the United States, USSR, France, 

Japan, and the United Kingdom (Finger et al., 1975). 

Comparison of temperature soundings found fair agreement (temperature 

difference less than *5C) among those of the United States, Japan, and 

the United Kingdom. A table of temperature adjustments has been recommended foruse with French and Russian data, which departed significantly 

from the other measurements. The wind measurements were in substantial agreement (root-mean-square differences less than 5 meters per second), 

except fox the Russian results above 40 kilometers. These tests, which 

involved 85 soundings in 29 comparisons, illustrate the difficulty in establishing a uniform reference for a global network. The accuracy of the 

rocketsonde data is estimated to be ±0.5 meter per second in velocity and 

I degree in direction for machine-calculated data (CIAP Monograph I, 

Chapter 9, 1975). Intercomparison measurements and comparison with 

cloud motion indicate a precision of approximately ±5 meters per sounding 

(Bauer, 1976). This error is too large for accurate determination of meridional flow, as pointed out by Danielson in CIAP Monograph I. Confidence in CMRN data will come at the expense of additional observation 

stations and more frequent soundings. International cooperation must be 

emphasized to achieve better uniformity in CMRN data. 

Realistic increases in the CMRN data rate are probably unlikely, and it is 

necessary to turn to satellite observations and other techniques to supplement the data. Objective analysis techniques for satellite radiation data 

and rocketsonde data exist (Gelman and Negatani, 1976), which can automatically prepare synoptic analyses of the 5-, 2-, and 0.4-millibar levels. 

The lack of vertical resolution of the satellite radiance information plus 

real space and time variability are the principal causes for the root-meansquare difference of 5 to 14 meters per second in velocities between rocketsonde and satellite data. The vertical height resolution is limited to about 

10 weighting functions representing the optimum infrared sounding. 

Vertical eddy-diffusion coefficients are important parameters that affect 

the lifetime of constituents in the stratosphere and their impact on ozone 

depletion. The current use of 1-D models to provide estimates of ozone 

depletion requires a realistic globally averaged eddy-diffusion coefficient. 

Vertical eddy-diffusion coefficients, Kz, are currently determined from 

measurements of altitude profiles of CH4 and NO. The derivation of 
MINOR SPECIES AND AEROSOLS 125 

Kz requires an accurate description of sources and sinks and the assumption that other forms of transport are not affecting the altitude profile. 

This latter assumption is seriously challenged below 20 kilometers by the 

effects of tropopause folding and cumulus convection in the tropics 

(Reiter, 1975; CIAP Monograph I). At the time of the NASstudy, only a 

few midlatitude profiles existed, which lead to an overconfidence in estimating the eddy-diffusion coefficients. Numerous N.0 measurements 

conducted by NOAA at tropical, middle, and high latitudes have been 

recently analyzed for eddy-diffusion coefficients (Schmeltekopf, 1977). 

Significant variations exist with latitude. Itis clear that location of the 

measurements is an important consideration. Further measurements are 

needed to cover expected seasonal variations. 

AEROSOL MEASUREMENTS 

The possible importance ,ofheterogeneous chemistry on aerosols in the 

stratosphere is discussed in the position paper by DeMore (1977). The 

situation with regard to chemical effects is characterized by a general lack 

of specific rate data for possible surface reactions. Fortunately, upperlimit arguments can be used to rule out known heterogeneous effects as 

being important for removal of CFM's, HC1, CIO, and Cl. However, 

DeMore (1977) concludes that possible surface catalytic processes involving hydrolysis of chlorine nitrate to produce HNO3 presently cannot be 

excluded as being competitive with gas-phase processes. Thus, aerosols 

may be important to certain aspects of the ozone depletion problem. 

An account of the various techniques used to measure aerosols may be 

found in the position paper of Rosen (1977). No single method can be 

used to make a definitive measurement of aerosols in the stratosphere. 

However, if the shortcomings of each technique are taken into account, 

there does seem to be a "mainstream" of agreement. 

a 	There is reasonably good agreement or consistency among the various methods of measuring the vertical profile of aerosols. The results obtained from simultaneous observations using a photoelectric 

counter and a lidar technique compare quite favorably. Details of 

the experiment and the model needed to make the comparison have 

been presented by Pinnick et al. (1976). Bigg (1975) has found good 

agreement between simultaneously measured vertical profiles as ob­
126 CHLOROFLUOROMETHANE WORKSHOP REPORT 

tained by the impactor technique and the photoelectric counter. 

Although the structure in the profiles may vary in time and space, 

there is essentially agreement that a layer exists at roughly 20 kilometers with a relative maximum in concentration and mixing ratio, 

This is to be contrasted with the profiles of condensation nuclei 

(CN, smaller particles) which show no relative maximum. In addition, there appears to be no consensus in the CN measurements 

above 20 kilometers. 

One of the most important aspects of the 20-kilometer layer is its 

variation with. time. It is generally more useful to study the mixing 

ratio at the maximum, because this quantity can only be increased 

by a source and atmospheric mixing will tend to decrease it. Thus, 

the variation of the peak mixing ratio with time is very dearly related to the source strength of stratospheric aerosols. 

The longest continuous record of a direct mixing-ratio measurement 

has been presented by Hofmann and Rosen (1977). The effect of a 

volcanic eruption in October 1974 is clearly evident. Although the 

aerosol decay was apparent after the volcanic eruption, it is not clear 

if the layer was in a steady-state (nondecaying) condition before this 

event. A typical average value of the particle concentration at 20 

ilometers for the latitude is about 1 per cubic centimeter for sizes 

larger than 0.3 micron in diameter. 

Although the time variation of smaller particles (CN) in the stratosphere has not been adequately studied, preliminary results indicate 

that it remains fairly constant (Rosen and Hofmann, 1977), again in 

contrast to the larger particles. 

Because of the enormity of the task, very few efforts have been 

made to systematically determine the global distribution of stratospheric aerosols. Some recent measurements have been reported by 

Rosen et al. (1975). These results appear to be fairly consistent 

with the global distribution of mass loading as reported by Lazrus 

and Gandrud (1974b), although above 20 kilometers there are very 

few mass measurements available for comparison. 

No systematic measurements have yet been reported concerning the 

global distribution of CN in the stratosphere. 
MINOR SPECIES AND AEROSOLS 127 

* 	 A summary of almost all direct measurements of the size distribution in the 20-kilometer layer has been presented by Harris and 

Rosen (1976). In spite ofthe many inconsistencies, there appears to 

be a mainstream of agreement. It is also apparent from the foregoing work that, generally, the impactor data give unrealistically 

high particle concentrations for radii larger than about 0.3 Micron. 

This effect can most likely be attributed to a fundamental problem 

in the technique itself: the inability to correctly sample volatile 

aerosols. When the suspect data are taken into account, the mainstream of agreement becomes even more pronounced. Calculations 

of the mass loading and lidar backscatter cross sections using absolute size distributions consistent with this mainstream of data give 

excellent agreement with the corresponding measured values (Pinnick et al., 1976). Thus, there is good reason to believe that current 

knowledge of the size distribution of stratospheric aerosols is reasonably adequate. 

" 	The major chemical component of the present aerosol in the 20­

kilometer layer is generally believed to be H2S4 . However, impactor samples indicate that one or more small insoluble dense nuclei 

can frequently be associated with each liquid particle. The source 

of these small nuclei could be terrestrial, extraterrestrial, or both. 

Areas of uncertainties in aerosol measurements are as follows: 

* 	It is popular belief that, except for after volcanic eruptions, the 20­

kilometer aerosol layer is m a steady-state configuration supported 

by sulfur-bearing gases diffusing up through the tropopause. The 

best direct data available do not clearly show the existence of a 

steady-state layer, and it is possible that the layer would not exist 

without volcanic injections. Until this question is settled by longterm surveillance employing a technique capable of obtaining good 

quantitative measurements, modeling of the stratospheric aerosol 

will be impossible. Without a realistic model, it will be difficult to 

address the effect of man's activity on this stratospheric constituent. 

* 	 Although the real part of the index of refraction can probably be 

determined accurately enough from the chemical composition, the 

imaginary part is relatively uncertain because it can depend heavily 

on the impurities present in the aerosol. The complete index of 

refraction is needed for optical models that assess the effect of 

aerosols on climate and relating remote sensing data to useful quantities. 
128 CHLOROFLUOROMEfHANE WORKSHOP REPORT 

* 	 The vertical profile of CN above 20 kilometers is still very uncertain. 

There is mounting evidence that an appreciable fraction of these 

particles in the stratosphere may be of extraterrestrial origin. Measurements of CN to at least 40 kilometers are needed to provide 

baseline information and asymptotic limits on the size distribution 

used in optical models. 

* 	 Although the overall size distribution derived from the mainstream 

of measurement does not appear to be incorrect, some thought 

should be certainly given to developing a definitive method for 

measuring this parameter because none of the present methods are 

entirely adequate in themselves. In addition, various types of direct 

measurements designed to test proposed optical models based on 

proposed size distributions should also be conducted. 

The major areas of uncertainty lie in the distribution of CN above 

20 kilometers and the existence or nonexistence of a steady-state 

20-kilometer aerosol layer. 

RECOMMENDATIONS 

MEASUREMENTS
 

Certainly, the present data base of both reactive and nonreactive stratospheric minor constituents needs to be improved and extended before the 

confidence in model predictions will be improved substantially. Although 

significant improvements in the data base have been made in the last few 

years, the stage is now set for -measurement programs that will provide 

considerably more stringent bounds for model simulations. 

A clear distinction should be made between the difference in measurement strategies for: (I) the source terms that are decoupled from the 

stratospheric photochemical scheme because of the long-time constants 

for photolysis or oxidation processes, and (2) the radical sink terms that 

are strongly coupled photochemically. For example, high priority should 

be given to simultaneous observations of HC1 and CIO to establish the 

free-to-stable chlorine ratio, whereas simultaneous measurements of CIO 

and CH 3 C1 are of minor importance. 

The proviso that extensive intercomparisons and cross-calibrations be 

made ufiderlies any continued measurement program. These must be 

made to settle existing outstanding discrepancies and to assign more 
MINOR SPECIES AND AEROSOLS 129 

meaningfully the accuracy of the present data base. In conjunction, sensitivity tests of the data inversion techniques of remote sensing methods 

should be undertaken. 

Although simultaneous measurements are of prime importance, dedicated 

exploratory measurements to identify new species or to confirm as yet 

undetected species are recommended. 

In addition to abundance measurements, programs should be developed to 

obtain in situ measurements of photodissociation coefficients and/or solar 

flux. At least relative measurements of direct and indirect flux would aid 

in the detailed model simulation of specific stratospheric missions. 

It is also of particular importance for laboratory measurements to be 

made, under simulated stratospheric conditions, of spectroscopic parameters necessary for the inversion of long-path absorption/emission data. 

Facilities are available for these studies, and use of them is encouraged. 

HYDROGEN SPECIES 

The data base of hydroxyl-profile measurements needs to be extended, 

and the complete altitude range needs to be covered. Hydroxyl measurements should be carried out with simultaneous measurements of at least 

03 and H2 0. Techniques should be developed for the measurement of 

H, HO2 , and H2 02 . Ground-based measurements of OH should be intercompared with simultaneous profile measurements to define the stratospheric contributions. Measurements of seasonal, temporal, and latitudinal 

variability of hydroxyl total amount and stratospheric concentrations 

should be made. First, 1-D models require at least hemispheric profile 

averages of abundance measurements. This necessitates a long-term experimental program to determine latitudinal, seasonal, and diurnal trends. 

To obtain reasonable hemispheric averages, it is recommended that several 

new, well-chosen launch sites, consistent with scientific objectives, be 

established to supplement present sites. An equatorial station should be 

given high priority. It is also recommended that experiments be conducted over relatively short periods of time during specific seasons to 

determine reliable seasonal trends. In this respect, greater priority should 

be given to these missions in order to better utilize existing launch stations. 

In the intervening periods and latitudes, ground- or aircraft-based column 

measurements are required to supplement the profile data and to permit 

routine monitoring as required to supplement the profile data and to 
130 CHLOROFLUOROMETHANE WORKSHOP REPORT 

permit routine monitoring of constituent variation. The chosen launch 

sit6s should maintain routine upper-air meteorological and ozone measurements in order to correlate measurements and to identify unusual atmospheric conditions. 

On a shorter time scale, and while fulfilling some of the requirements of 

the long-term studies, more extensive programs designed for critically 

testing the accuracy of model chemistry by simultaneous diurnal and profile measurements are required. Simultaneous measurements of interrelated components of the HOx, NOx, and CIOX cycles should be given 

greater priority. With present instrumentation, collaboration of several 

groups would enable key experiments to be attempted in a relatively short 

time. To define and organize these missions, the establishment of a panel 

of active scientists is recommended. 

Isolated measurements of OH are of little value because it has been demonstrated that this species is apparently briefly variable, and this requires 

careful definition of its chemical environment (i.e, H2 0, 03, 0, and T). 

It is also of little value to measure C1O and HCl or NO, NO2 , and HNO3 

without knowledge of the OH density, because the (ClO)-to-(HC1) ratio is 

directly proportional to (OH), and the (NO2 )/(HNO3 ) ratio is inversely 

proportional to (OH). 

NITROGEN SPECIES 

Although reasonable attempts have been made to measure the ratio of 

species that are expected to be photochemically controlled, none of the 

NOXINO ratios so determined are really in satisfactory agreement with 

specific model calculations. Further studies by all in situ or all long-path 

techniques are required, as relative measurements of higher accuracy can 

be made. In view of the current strong CIOx-N0 X coupling, simultaneous 

missions of C1Ox, NO., and HO instrumentation should be attempted. 

Simultaneous diurnal NOx measurements using several techniques should 

be undertaken to clarify the longstanding controversy regarding the behavior of NO 2. 

The observations that total NO2 (and thus presumably NO) essentially 

disappears north of 50 in winterhave yet to be explained quantitatively. 

Both multidimensional model calculations and winter profile measurements at high latitudes are required. 
MINOR SPECIES AND AEROSOLS 131 

The role of N.O. in stratospheric chemistry needs further clarification by 

measurements to determine profiles, especially in winter at northern and 

southern latitudes. This is important because N2

0 S acts like a reservoir 

for inactive odd nitrogen. 

A detailed assessment of the influence of the rapid N0 2/NO conversion 

on long-path-measurements at twilight is required. 

HALOGEN SPECIES 

The ratio of CI0 to bound forms of chlorine is critical to the depletion of 

ozone by CFM. New techniques should be developed to measure CIO in 

order to support the resonance fluorescence measurements. Simultaneous 

measurements of CIO, Cl, HCO, 03, OH, and CH4 are required to test current chlorine photochemistry models. Measurements of CIONO 2 , in conjunction with C10 and HC, are urgently required. Initial measurements of 

C1ONO 2 should be made at dawn because concentrations will be larger. 

New measurement techniques, such as resonance fluorescence, should be 

developed for HC1. The latitudinal and seasonal variations of HCI and CIO 

should be measured. 

Other possible chlorine compounds should be investigated, and attempts 

should be made to measure them if they are determined to be important 

to atmospheric measures. Filter-impregnated techniques for measuring 

eltorine compounds should be developed further. The development of 

any technique to measure the total chlorine content would be a major 

advance. Simultaneous measurements of HC1 and HF should be made, and 

attempts to develop techniques for measuring other possible fluorine compounds, such as COF2 and COFC1, should be pursued. A possible technique is the further improvement of high-resolution spectroscopic instrumentation. 

Concentration proffles of HBr and inventories should be measured. 

The injection of chlorine, fluorine, and sulfur compounds by volcanoes 

into the stratosphere should be monitored. 
N77a31698
 

CHAPTER 4 

ONE-DIMENSIONAL MODELS 

INTRODUCTION 

The one-dimensional (1-D) model is a particular simplification of the 

atmospheric radiative/chemical/dynamical system in which chemistry is 

treated in greater detail than radiation or dynamics. Although the amount 

of chemistry in a I-D model varies from model to model depending on 

the problem being examined, in general, it is as complete as our knowledge 

of laboratorykinetics permits for the simulation of the atmosphere. Transport is introduced parametrically in the diffusion approximation as the 

vertical altitude-dependent eddy-diffusion coefficient, K. Radiation in 

the ultraviolet and visible wavelength range is treated in some detail, 

sometimes including the effects of molecular multiple scattering and surface or cloud albedo. Radiative feedback through changes in infrared 

heating or cooling and changes in ultraviolet heating have been examined 

but are not usually included in 1]) chemistry models at this time. A more 

complete description of these processes isincluded in Chapter 5. 

Because the effect of the chlorofluoromethanes (CFM's) on ozone is primarily through chemical processes, the 1-D chemistry model has become 

the primary predictive tool for assessing possible changes in the natural 

ozone content. There are many caveats on the use of l-D models. These 

have been discussed in the Climatic Impact Assessment Program (ClAP) 

Report of Findings, as well as in the National Academy of Sciences (NAS) 

reports on stratospheric flight and on CFM's. Chapter 5 of this report discusses many of the modifications that may have to be made to 1-D model 

results to account for the effect of feedbacks through radiation and dynamics not included in the models. 

The focus of this chapter will be the current status of predictions by 1-D 

models. The first three sections discuss some aspects of what might be 

called model mechanics: the parametric simulation of vertical transport, 

the effects of diurnally varying reactants, and molecular multiple scattering. This is followed by three 'sections on the status of understanding 

of stratospheric measurements in terms of current theory. Next, the 

question of possible sinks for CFM's isdiscussed, followed by a discussion 

of the uncertainties involved in making predictions of stratospheric 

ozone change. 

!Pre; 'djng, page blank! 
134 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Special care should be taken in interpreting the subsections on comparison of theory and measurements. In general, the comparison is between 

the result of a 1-D model and a measurement made at one point in space 

and time. The interpretation of a l-D model calculation is still a subject of 

discussion. In general, it represents a global average at each altitude of the 

species being calculated. By comparison with data, this global average has 

been shown to resemble the situation at around 30'N latitude (J6hnston 

et al., 1976). Because most of the measurements of minor constituents 

have been made at this point, there is reason to expect gross quantitative 

agreement between the measurements and model results. The degree of 

agreement expected depends critically on what is being compared. For 

instance, when concentration ratios of species that are closely linked by 

fast chemical reactions are compared, agreement should be expected within the uncertainties of laboratory data. Good agreement is more likely 

when the concentrations of the controlling long-lifetime species are wellknown. When these supporting measurements are not made, an uncertainty in the comparison comparable to the expected variability of the particular species must be included. When a reactive species is measured without any supporting measurements, the interpretation must include the 

effects of the temporal and spatial variations arising from a species' relationship with all of the other atmospheric constituents. Such variations 

might limit the expected agreement with a 1-D model to within a factor 

of 2. 

PARAMETRIC SIMULATION OF NET VERTICAL TRANSPORT IN 1-D 

MODELS 

Net vertical-transport processes in 1-D models of the stratosphere are 

usually represented through diffusion approximations. The net verticaldiffusion flux, Fz, ofany minor trace species in the stratosphere is assumed 

to be proportional to the gradient of the mixing ratio of that trace species: 

Fz =- pK -y (C./p) (4-1) 

where C. is the concentration of the ith trace species, p is the air density, 

and z is the altitude. The diffusion coefficient, KZ, is usually called the 

vertical eddy-diffusion coefficient. In both the conceptual and operational 

sense, the term eddy diffusion in 1-D models is a misnomer. In the past, 

this use has led to unnecessary confusion in the interpretation of l-D 

models. Consequently, in the following discussion, Kz will be referred to 

as the I-D vertical-diffusion coefficient. 

41, 'JAY. ~ Il 
ONE-DIMENSIONAL MODELS 135 

With this representation for the net vertical flux, the time-dependent 1-D 

continuity equation for the species concentration becomes 

at 7K )]i+Pi at az z [PXaz (OikP - L (4-2) 

where Pi and LI are the photochemical production and loss terms, 

respectively. When appropriate, a third term, Si, representing any other 

anthropogenic sources or sinks of Ci are also added to the right side of 

equation 4-2. If the concentration of C_ is at steady state, equation 4-2 

becomes 

z pK -z (C/P)] + Pi - Li = 0 (4-3) 

There is a large degree of uncertainty in the interpretation of equation 

4-2 or 4-3. Ideally, with both proper time and space averaging of the 

appropriate variables, K,, P., and Li, these equations can at least be considered a global representation- Under this ideal condition, the remaining 

question would be the yet to be demonstrated suitability or accuracy of 

the diffusion approximation, equation 4-1. Unfortunately, this ideal condition does not exist. Because of the temporal and spatial variations of 

tropopause height and the intrinsic differences in the dynamic processes 

of the troposphere and stratosphere, there is no apparent technique for 

averaging K., PV1 and L. satisfactorily in the global sense. Furthermore, 

such averaging would require detailed knowledge of the spatial distribution 

of the trace species at least qualitatively, if not quantitatively. At the 

present time, such detailed information is mostly unavailable. Consequently, the currently used values of Kz derived from limited data are somewhat 

different. As will be discussed below, these differences are real and must 

be considered as part of the uncertainties in l-D stratospheric models. 

CONCEPTUAL BASIS FOR THE INTRODUCTION OF Kz 

Since primary concern is with the minor trace species of the atmosphere, it 

is reasonable to assume that the dominant motion is not dependent on 
136 CHLOROFLUOROMETHANE WORKSHOP REPORT 

the local kinematics of individual trace species. For any particular 

species, Ci, if equation 4-2 or 4-3 is applicable and if the loss and production processes are known, then, given the measured vertical profiles of C., 

equation 4-2 or 4-3 can be inverted to obtain K . If all the assumptions 

noted previously are correct, then K. so obtained should be independent 

of particular choices of tracer species. Based on current understanding of 

this kind of analysis, it seems reasonable to ascribe the observed K. variations to general incompleteness of data rather than to gross error in the 

conceptual assumptions noted above. 

Two kinds of atmospheric tracers have been used to derive K. vertical 

profiles. First, there are the chemically reactive tracers, such as CH4 and 

N20, that are mostly natural in origin. The other tracers are chemically 

inert; namely, the radionuclides introduced into the stratosphere by past 

atmospheric nuclear tests. In addition to these tracers, numerical experimeats can also be conducted on existing three-dimensional (3-1D) general 

circulation models (GCM's) with idealized inert tracers. By averaging the 

3-D results from these numerical experiments, an effective 1-D verticaldiffusion coefficient can also be derived. The results from each of the 

three types of tracer analyses will be summarized and compared. 

CHEMICAL TRACERS 

CH 4 and N2 0 are identified as reasonable tracers because of their relatively simple chemical-loss mechanisms. Wofsy and McElroy (1973) first 

derived a K. proffie from the available CH4 data. Later, Hunten (1975) 

used the same chemistry (CH 4 loss rate) with a different smoothing of the 

CH4 data to obtain a somewhat different K profile. Recently, Dickinson 

(1976), using both CH4 and N20 data and the loss rates as given by Chang 

(1976), has carefully analyzed the various derived K. profiles arising from 

differences in data interpretation. Profile B in figure 18 represents an 

average of Dickinson's results. Profile A is qualitatively similar to the 

result of Hunten (1975), but multiplied by a factor of 2. This last modification was necessary because the original Hunten profile did not produce 

a chemical loss rate of CH4 that is consistent with that used in its derivation. As can be seen, these two profiles (A (NASA, 1 9 7 6 t) and B (Chang, 

1976)) are very similar above 25 kilometers. The differences between 10 

and 25 kilometers will be discussed in the next section. 

*R. S. Stolarski, private communication to model groups giving parameters for model 

comparison. 
60 

50 

--A 

-A 

40 

Az 

o 

~ 30 

00 

10 3 •104 ,cm1C 105 so,, 

Figure 13. Vertical-diffusion coefficient profiles. 
138 CHLOROFLUOROMETHANE WORKSHOP REPORT 

These tracers are useful in that they should be fairly uniform globally and 

that more than any other tracers they may be assumed to be in steady 

state. However, their apparently simple chemical loss mechanisms are 

quite deceptive. The principal loss mechanism for CH 4 is its reaction with 

OH, whereas for N.0 it is by photodissocation. The concentrations of 

OH and the photodissociation coefficients JN 0 are very model-dependent 

and contribute significantly to the uncertainties in the derived Kz . 

Another source of uncertainty is the lack of data above 35 kilometers for 

both CH 4 and N.O. There is only one set of rocket data, which, in view 

of natural variability of all atmospheric trace species, would have to be 

construed as insufficient. There is also a need for more global data coverage for CH 4 , because most of the existing data in the stratosphere are 

from one location (Ehhalt, 1972 and 1974); It is unlikely that these uncertainties can be removed in the near future except, perhaps, by the 

gathering of more high-altitude data. 

RADIONUCLIDE TRACERS 

Past atmospheric nuclear weapons tests have injected an easily identifiable 

amount of radionuclides into the stratosphere. These radionuclides are 

, Rh 102 , Cd 1 0 injected either as fine particles (e.g., Sr 9 ° , W' 8 5 , and Zr 9 s) 

or as a true gas (C' 4 ). Because there are no chemical sources or sinks 

(except for well-understood radioactive decay) for these tracers, the inversion of equation 4-2 is particularly simple. Unfortunately, this advantage 

over chemical tracers is almost totally overshadowed by the difficulties 

caused by the transient (time-dependent) nature of its removal from the 

stratosphere. Uncertainties in sampling initial distributions and all sftbsequent distributions in the stratosphere and troposphere raise serious 

doubts about the conclusiveness of this type of analysis. A detailed discussion of these problems is in Chang (1975). Nevertheless, the available 

results from the particulate data seem to agree with profile B in the altitude range 10 to 25 kilometers. However, C14 data agree better with 

profile A divided by 2 (Johnston et al., 1976). As a result, there is at least 

a factor of 2 uncertainty in using these tracers. These results, coupled with 

the requirements from stratospheric modeling of CH 4 and N.O , necessitated the present acceptance of both profiles A and B as possible representations in 1-D model. Fortunately, although profiles A and B gave different results for transient behavior, steady-state calculations using these 

two profiles yield identical predictions -of the effect of CFM's on ozone 

(Chang and Wuebbles, private communication, 1977). 
ONE-DIMENSIONAL MODELS 139 

GCM TRACER EXPERIMENTS 

Mahlman (1975) conducted two numerical experiments using the transport 

coefficients from the _GCM at Geophysical Fluid Dynamics Laboratory/ 

National Oceanic and Atmospheric Administration (NOAA). One was a 

midlatitude point-source injection experiment, and the other was a 

stratified tracer experiment. Using the time history of tracer distributions 

from these experiments, he was able to derive two profiles of effective 1-) 

vertical-diffusion coefficients that turn out to be in close agreement with 

profile B in the altitude range of 10 to 20 kilometers. Between 20 and 25 

kilometers, the results from the experiment differ by a factor of 10, but 

both profiles A and B fall in the middle of his values (more details can be 

found in his paper). Again, because of the transient nature of Mahliman's 

(1975) numerical experiments and the lack of comparative analysis from 

other models, this agreement is not sufficient to demonstrate a preference. 

The important point in the assessment of the l-D vertical-diffusion coefficient, Kz,is the general agreement between the average values of profiles A and B in the range 10 to 20 kilometers. For the SST-NOX 

problem, the details of Kz profiles in this altitude range are critical because 

this is the region of NO, injection. However, for the CFMCl x problem, 

the average effective transport rate across this region is the important 

parameter because CFM's from the ground must pass through this region 

to higher altitudes to be effective in releasing C1 to catalyze the destruction 

of ozone. Because the average transport rate across this region is similar 

in this respect, the effect should also be the same at steady state, as was 

found by Chang and Wuebbles (private communication, 1977). Consequently, the sensitivity of the ozone distribution to K is lower than 

might at first be expected. 

DIURNAL AVERAGING EFFECTS 

For an atmosphere with constituents that have chemical or dynamical 

lifetimes comparable to or less than 24 hours, the natural time-dependent 

effects must be carefully considered. The direct approach of integrating 

the diffusion equations with a time step short enough to properly reproduce the essential diurnal effects is too expensive because of the approximately 3 years of model time necessary for achieving true cyclic 

behavior from a steady-state starting point. Because of this fact and the 

need to follow the effects of atmospheric perturbations for a century or 

longer, some form of diurrnal averagingis necessary. Diurnal averaging is 

a procedure that permits the photolysis and gas kinetic rates to be modified 

so that the long-term behavior of a steady Sun integration is the same as 
140 CHLOROFLUOROMETHANE WORKSHOP REPORT 

the true diurnal average of the long-time integration of the fully timedependent calcuation. 

A typical example of a noxlinearterm contributing to the destruction rate 

of odd oxygen is 

k1 [0] [CIO] (cm- sec -

l) 

corresponding to the reaction 0 + CIO - 21 + 02 with the reaction rate 

constant k1 . The diurnal integral of this rate is represented by 

24 hr 

-f k 1 [0] [CIO] dt 

0 

and is frequently approximated by the simple form 

RA-Ik [0] [1] T 

where 

T = 86,400 seconds 

To evaluate the diurnal effects on ozone reduction due to CFM's in the 

presence of C1ON0 2 , Liu (1977) has run a steady-state program with 

24-hour averaged photolysis rates for 300 N equinoctial conditions. The 

model used a constant 1974 CFM injection rate resulting in a steady-state 

CIx mixing ratio of 5.67 parts per billion (ppb). The reaction rate of 

CIONO formation was 

- 1.53 X I0 14 exp (1O00/T) 

117 X l0Ie8 exp (250/T) + M 

In steady state, this run gave a column-ozone reduction of 3.5 percent. 

When the model is run in the actual diurnal mode, the value of R is found 

to be 60 percent larger than RA at an altitude of 30 kilometers. Because 
ONE-DIMENSIONAL MODELS 141 

of this type of result and the prohibitive cost of direct-time integration for 

large net model times, care must be taken with the method of diurnal 

averaging. 

Lin's (1977) approach is to make use of the fact that diurnal effects on the 

long-lived species are small. This permits their densities to be approximated by the solutions of the steady-state model (fixed Sun). For the 

present calculations, the profiles of H2 0, CH 4 , H., NO CF2C2' CFCI3 , , 

CO, NOx, and C1, are held fixed at their steady-state values, whereas Ox, 

HOX, H2O2, O(lD),O,03, H, OH, H0 2, N NoNOV HN03 NOs' N2 050, lCIO, , HCI, and CIONO 2 are allowed to vary with time. In addition, for the cases shown in figures 19 through 22, only Ox, HOx, and 

H20 2 are calculated with diffusive flow. 

Let the subscripts denote the steady-state result and d, the diurnal result. 

Then, after 120 days of integration in the diurnal model, two examples of 

diurnally averaged rates are: 

12 hours 24 hours 

fOld[NO 2 ld dt = 0.74f [0 1s [NO 2]s dt 

and 

12 hours 24 hours 

f od[aO]d dt = 1.6fo [0] [CO] s dt 

0 0 

where subscripts d and s denote diurnal and steady-state values, 

respectively. These results are shown in figures 19 and 20, where the 

24-hour averages of 2 [0] [NO2 ] and 2 [0] [CIO] in the diurnal model 

are compared to the steady-state model showing that, at 30 kilometers, the 

C1x perturbation due to the introduction of CFM's is enhanced by a factor 

1.6 from the CIO diurnal effect alone. The factor is 1 without CIONO 2 

and increases with increasing formation rate of CIONO 2 . The foregoing 

calculation also indicates that the ozone reduction due to the NOx perturbation is reduced by 25 percent because of the diurnal effects on NO2 

alone. Even though the test cases were not run to a converged cyclic condition, the trend indicates that the final ozone reduction for a true diurnal 
60I 1111111 I II i III I 1111II 

f2 24 hr 

-o 

30 

20 

1011 

C 

Fgure 19. 

10 10 

LOSS RATE PER DAY (p0cm 3 -day) 

Effet of diuralaveraging on the 0 + N0 2 

10 11 

reaction rate, 
60 1 I1 1 1 1 1 ' I 1 1 1 1 1 1 

f 

24 hr 0IC dt 

so 

40 -­ n 

30 z 

-~ 40 

20 

100 

10o I I 111 I I 1 1 1 1 1 1 11 I1 I 1 1 1 1 1 1I 

1010 1011 

LOSS RATE PER DAY (#i cm3 -day) 

Figure 20. Effect of diurnal averaging on the 0 + CIO reaction rate. 
144 CHLOROFLUOROMETHANE WORKSHOP REPORT 

model is probably about twice the ozone reduction predicted by the 

steady-state model. 

Figure 21 shows the diurnal variations of CIO along with the corresponding 

steady-state values,which are indicated by arrows. The variations below 40 

kilometers are from the conversion of CIO to C1ON0 2 at night and the reappearance of C10 in the morning by photodissociation of ClONO2 . Without C1ON0 2, 0IO behaves as expected and has no diurnal variation below 40 

kilometers. Below 50 kilometers,the variation of CIO isdue to Cl + 03 at 

night and its destruction by CIO + 0 during the day. The-diumal variations of ClON0 2 and its steady-state values are shown in figure 22. The 

daytime CIONO2 density is about 20 percent less than the steady-state 

values between 20 and 30 km. From the foregoing discussion, it is clear 

that the problem of diurnal averaging is quite important in models that 

attempt to calculate the net ozone reduction in the presence of CIONO2 . 

MULTIPLE-SCATTERING EFFECTS 

The effects of including multiple scattering in the photodissociation-rate 

calculation are of three types: (1) the effect on photodissociation coefficients, (2) the effect on ambient species concentration profiles, and (3) 

the effect on model sensitivity to perturbations. -The results shown in 

the following subsections were obtained using the Lawrence Livermore 

Laboratory 1-D transport-kinetic model (Luther and Wuebbles, 1976), 

which includes C1ONO 2. The effect of multiple scattering was incor­ 'porated into the photodissociation-rate calculation by applying correction 

factors obtained from a highly detailed solar-radiation model to the flux, 

F, based on a pure absorption calculation. A solar zenith angle of 450 was 

used, and a surface albedo of 0.25 was assumed. The results are qualitatively 'representative of multiple scattering calculations made using other 

models. 

PHOTODISSOCIATION COEFFICIENTS 

Because virtually no radiation reaches the lower troposphere at shorter 

wavelengths, the effect of multiple scattering is largest for species having 

strong absorption cross sections at wavelengths greater than 290 nanometers. Computed photodissociation coefficients, including multiple 

scattering (TMs) and assuming pure absorption (JpA), are compared in 

table 28 for the same ambient atmosphere. 
ONE-DIMENSIONAL MODELS 145 

30 km 

30 km 

40 Im 

40 km 

107 108 50 km - N 

-- 50 km 

° 	 4 8 12 20 	 24 

LOCAL TIME 1hr) 

12 	 16 

Figure 21. Diurnal variations of CIO along with 

the corresponding steady-state values. 
146 

109 

108 

CHLOROFLUOROMETHANE WORKSHOP REPORT 

-

0kkm 

30kmn 

40 km 

Cx 

10 72
 

4 812
 12108 16 20 24 


LOCALTIME (hr) 

Figure 22. Diurnal variations of CIONO 2 along with 

the corresponding steady-state values. 
ONE-DIMENSIONAL MODELS 147 

Table 28
 

Comparison of Photodissociation Rates Calculated With and
 

Without Multiple Scattering (A = 0.25) 

Altitude (kilometers) 

03 +h -+ o(ap) + 02 

10 


20 


30 


40 


0 3 + h10 


20 


30 


40 


O(D +0)+02
 

NO2 + hI- NO +0 

10 


20 


30 


40 


HNO +hv - OH +NO 2
 

10 


20 


30 


40 


H2 0 2 +h - 2H 

10 


20 


30 


40 


CIO + h C1 

10 


20 


30 


40 


CIONO 2 + hv ­

10 


20 


30 


40 


0 

CIO+ NO 2 

JPA 


202X 10'4 


2.07 X 10-4 


2.41 X 10-4 


2.74X 10-4 


5.3X 106 


7.97X 10.6 


-5
 5.78 X 10


932X 104 


4.72X 10-3 


4.74X 10-3 


" 4.85X 10 3 


103
 4.97 X 


2.72X 10-7 


337X 10-7 


6
 4.98 X 10.


3 52X 10-s 


1.90X 10-6 


2.23X 10.6 


6-57X 10-6 


.351X 10-s 


1.72X 10-5 


2.55X 10-5 


1.92X 104 


1.64 X 10-3 


3-31 X 10-5 


3 43 X 10-5 


-
 5.22 X 10


2.24 X 10-4
 

0-4
 3.01 X 


3_00 X 10-4 


3.21 X 1014 


3.49X 10-4 


1.08X 105 


-5
 1.22X 10
-
 627 X 10


.4
 9.14X 10


7.77X 10-3 


7.76X 10-3 


7.68X 10-3 


7.73 X 10-3 


482X 10-7 


5.18X ]0-7 


5.11 X 10 6 


3.52X 105 


3.35)X 10.6 

3.46X 10-6 


7.29X 10-6 


"
 5
 3.52X 10
 

105
 3.25X 

3.82X 10-s 

"4
 2.05X i0 .
 1.62 X 103 


-
 5.57X 10


5
 5.61 	X 10'
" 
 7.07 X 10
 

2 38 X 10-4
 

JMS'JPA 

1.49 

1.45 

133
 

1.27 

1.85 

153
 

1.08
 

0.98 

1.65 

1 64
 

1.58 

1.56 

1.77
 

1.54 

1.03 

1.00 

1.76 

1.55 

1.11 

1.00 

189
 

1.50
 

1 07 

0.99 

1.68 

164
 

1.35 

1.06
 
148 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Only those photodissociation reactions significantly affected by multiple 

scattering are shown in table 28. In an attempt to account for the diurnal 

variation of photodissociation in the model, the solar flux was halved. 

The importance of the significant changes in photodissociation-rate coefficients is reflected in the species concentration profiles and in the model 

sensitivity. 

SPECIES CONCENTRATION PROFILES 

The effects of multiple scattering on the individual species concentrations 

making up Ox, HOx, and NOx are shown in figures 23, 24, and 25, respectively. The large percentage increase in O(1D) near 10 kilometers shown 

in figure 23 occurs when the reference concentration is very small; nevertheless, it is significant in terms of stratospheric chemistry. The increases 

in O(3P) and O(1D) in the 20- to 30-kilometer region are due to increased 

photolysis of 03. Because of differences in reference concentrations, a 

small percentage decrease in 03 causes large percentage-increases in the 

other species. The increase in 0( P) near 40 kilometers is mainly caused' 

by increased photolysis of NO2* The increase in 03 at this height occurs 

because the chemical production of 03 by 0 + 0i is increased more than 

the photolysis of 03. 

As 02S 

40 

3 0p) 

30 

20 

-20 0 20 40 60 80 

CHANGE IN CONCENTRATION tpert-nt) 

Figure 23. Change in O× species concentrations attributable 

to multiple scattering with A. = 0.25. 

10 
149 

so 

40 

- 30 

Figure 24. 

10 

40 


30 

20 

20 ­

-40 

Figure 25. 

ONE-DIMENSIONAL MODELS 

I I 

A, 025 

22 

-20 0 20 40 60
 

CHANGE IN CONCENTRAITION (eruent)
 

Change,in HOx species concentrations attributable 

to multiple scattering with As -- 0.25. 

H 

O N 

~A, - 0.26 N 

A II .. I I

2 

I 0 

0 40 0 120 160 

CHANGE IN CONCENTRATION (perfet)

Change in NOx species concentrations attributable 

to multiple scattering with As = 0.25. 
150 CHLOROFLUOROMETHANE WORKSHOP REPORT 

As shown in figure 24, the increase in OH results from HNO 3+ hv -) OH + 

NO2 and H20 2 + hp'-* 20H. Because the peak concentration in H 202 

occurs near 28 kilometers, a small percentage increase in H2 0 2 in this 

region can cause a large percentage increase in OH. The increase in HO2 is 

chemically linked to the increase in OH. 

Figure 25 illustrates the very large increase in NO near 20 kilometers due 

to increased photolysis of NO . The NO2 concentration increases because 

of increased photolysis of HNG3 . There. is very little HN0 3 above 30 

kilometers, so that [NO 2] decreases in this region result in [NO] increases, 

because it is essentially the only source of NO. As figure 25 indicates, the 

[NO] I[N0 2 ] ratio is changed significantly in the region between about 

10 and 30 kilometers. 

As shown in figure 26, the concentration of CIONO2 is reduced 20 to 40 

percent between 20 and 30 kilometers, which is the region of maximum 

CIONO2 concentration. Photolysis of CONO2 affects several other chlorine-containing species through a complex chain of reactions. Some of the 

key reactions are: 

CIONO2 + hp -+ CIO + N0 2 

CIO + NO ->C + NO 2 

Cl + CH4 -* HCI + CH 3 

Because -the peak concentration of CIONO 2 occurs near 25 kilometers, 

photolysis of CIONO2 acts as a source of CIO in this region. The large increase in NO between 20 and 30 kilometers tends to destroy C10, leading 

to a net decrease in CIO in this region and an increase in Cl. The increase 

in Cl leads to an increase in HC1 (figure 27) through the reaction with CH4. 

These results also indicate that inclusion of C1ONO2 has a more significant 

effect on C x species thanthe choice of surface albedo. 

The significant changes in the species concentration proffles are reflected 

in the total-column abundances. Ratios of total-column abundances computed with multiple scattering to that with pure absorption are 0.94 (03), 

0.975 (NO2), and 1.21 (NO). 

Although the foregoing analysis is overly simplified considering the complexity and interaction of the various chemical cycles, the results and discussion are an attempt to summarize the major mechanisms by which 

multiple scattering changes the species concentrations. 
ONE-DIMENSIONAL MODELS 151 

s0 / -

CIONO 2 

A. 0 

20 

10 

0 * I *, 

-00 40 -20 0 20 

CHANGE IN CONCENTRATION (pecent 

Figure 26. Change in concentration of CIONO 2 attributable 

to multiple scattering for A. = 0 and 0.25. 

HCI 

40 40 As=O A =025 

A,­

0 

WITHOUT CIONO 2 

0 I 

0 10 20 30 40 50 

CHANGE IN CONCENTRATION (pecent) 

Figure 27. Change in concentration of HCI 

attributable to multiple scattering. 
152 CHLOROFLUOROMETHANE WORKSHOP REPORT 

MODEL SENSITIVITY 

Model sensitivity to the release of CFM's at constant production levels 

was tested for cases-with and without CIONO2. The CFM calculation to 

steady state assumes that CFC 3 and CF 2 C 2 are released at 1973 rates as 

estimated by McCarthy (1974). 

Multiple scattering significantly reduces the sensitivity of the model without CIONO 2. When multiple scattering was included, the ozone reduction 

decreased to about 0.8 ofthat computed without multiple scattering. This 

reduction factor is nearly independent of surface albedo. The model with 

CIONO 2 is the more realistic model according to the present understanding. Essentially no change occurred in model sensitivity when multiple 

scattering was included in the model with ClONO 2. 

The decrease in sensitivity for the model without C1ONO2 occurs because 

multiple scattering increases NO 2 photolysis, which increases the NO concentration. Consequently, the importance of the reaction, CIO + NO, is 

increased relative to CIO + 0, thus decreasing the effectiveness of the 

chlorine catalytic cycle. MultipI6 scattering causes inedeased ClONO2 

photolysis and a reduction in the C1ONO2 coneentration: Therefore, 

although the ability of-the C10 x 
catalytic cycle to destroy 03 isreduced

by multiple scattering, the ameliorating effect ofC1ONO. is also reduced. 

SUMMARY
 

The effect of including multiple scattering in the photodissociation-rate 

calculation is to significantly alter the photodissociation rates and species 

concentration profiles. This is particularly true for those processes that 

are sensitive to wavelengths greater than 290 nanometers. The effect of 

multiple scattering on the model sensitivity was small for the model including CIONO2 . The small change in sensitivity occurs in spite of the fact 

that multiple scattering lessens the effectiveness of CIONO 2 in interrupting 

the 0 3-destroying catalytic cycles. 

COMPARISON OF THEORY ANDMEASUREMENT: HOX 

Among the major subgroups of atmospheric trace constituents that have 

been identified, the family of hydrogen atom radicals-H, OH, HO 2, and 

their close relative, H2,O-is probably the least understood. For example,

only OH has ever been measured in the stratosphere. Although the 
ONE-DIMENSIONAL MODELS 153 

infrared emissions of vibrationally excited hydroxyl molecules have been 

extensively monitored in the past, these data have not led to reliable determinations of H concentrations. 

The concentrations ofhydrogen gases are strongly affected by the distribution and variability ofthe hydrogen atom source compounds-water vapor, 

molecular hydrogen, and methane. These species have been discussed at 

length in both the CLAP and NAS reports. For example, enough watervapor measurements have been made to verify its extreme variability in the 

lower stratosphere, and HOx concentrations will reflect this variability to 

a large extent. In discussing hydrogen-oxide measurements, this potential 

source of uncertainty should be remembered. 

Figure 28 summarizes the OH data that have been collected to date. All 

but one of the OH measurements have been reported in the past 2 years, 

with many of these experiments using similar resonance fluorescence techniques. Also shown in figure 28 are two theoretical curves representing 

the results of typical model calculations both with and without the new 

rate constant for HO2 + NO (i.e., 8 X 10-12, see reaction 31 in table 1). The 

theoretical profiles were obtained using a diurnal averaging procedure and 

indicate average daytime values for OH. When comparing such data with 

theoretical calculations,, one significant point to consider is that the stratospheric and mesospheric OH abundances presented by Anderson (1971 

and 1976) were obtained at large zenith angles so as to reduce sunlight 

interference in the experiments. (Anderson's observations were at 860 in 

1971 and 800 in 1976.) OH concentrations will decrease rapidly.in twi-' 

light. The problem of extrapolating the OH measurements at large zenith 

angles to average daytime concentrations involves the use of diurnal 

models such as those previously described in this chapter. The extrapolation factor is probably between 1.5 and 3. However, the determination of 

this factor is based on some of the less certain chemical-rate constants. 

Scaling Anderson's (1971 and 1976) stratospheric OH concentrations upward by a factor of 2 throughout the stratosphere and mesosphere indicates that they are roughly compatible with the lower end ofthe range of 

the determination by Burnett (1976) of total integrated vertical daytime 

OH column abundance (molecules cm-2 ) (see Chapter 3). 

The OH airglow data indicate that mesospheric OH, formed mainly by the 

reaction H + 0 , exhibits a large variability, with a time scale of hours to 

days. Whether or not this is the major source of variability in the Burnett 

(1976) data is unresolved. Also, it is not known how much of the OH 

column variability occurs in the stratosphere. 
154 CHLOROFLUOROMETHANE WORKSHOP REPORT 

70 

60 

noi....--­

50 ST-TOAUE 

* 40 

0 

20 

TROPOPAUSE --i 

10 

01 Ilii 

10 10
 

OH CONCENTRATION (molecules CM3

)
 

Figure 28. Observed atmospheric hydroxyl radical concentrations. Superimposed are two theoretical curves with (- ) and without ( ----) the 

new rate coefficient of 8 X 10-12 cm 3 /sec for the reaction of NO + HO2 

(reaction 31 in table 1). Caution must be exercised in comparing calculated average OH concentrations in daylight with atmospheric observations at large zenith angles. 

There is a significant gap in the OH data between 10 and 30 kilometers. 

In these altitude regions, OH partially controls the rate of conversion of 

HC1 into CIO and, thence, into chlorine nitrate. 

Early modeling efforts were hampered by the virtually complete lack of 

knowledge of OH abundances in the stratosphere and the concurrent lack 

of knowledge of most of its important reaction rates. Anderson's (1971 

and 1976) stratospheric OH measurements improved this situation by indicating the expected magnitude of [OH]. The agreement in the shape of 

the experimental [OH] altitude profile with calculated ones encouraged 
ONE-DIMENSIONAL MODELS 155 

Donahue et al. (1976), among others, to use these measurements in an attempt to narrow the choices for the critical reaction rates OH + HO 2 and 

O + H0 2 . The rapidly improving laboratory reaction-rate data for HO2 is 

beginning to permit deduction of OH concentration from kinetic data. 

Reasonable results for the magnitude of [OH] can be obtained without 

arbitrary adjustments of the model parameters. It isdifficult to assess the 

validity of such agreement because the present data base is so meager. 

The recently measured rate coefficient fox HO2 + NO -> OH + NO2 of 

8 X 1012 cm3!see, is a factor of 40 greater than that previously used at 

stratospheric temperatures. As can be seen from figure 29, the new HO2 + 

NO rate constant systematically increases the calculated OH concentration 

at all stratospheric altitudes, but particularly so below 40 kilometers. For 

example, at 30 kilometers, the OH increase is a factor between 1.5 and 2, 

depending on the specific model employed. 

The acceleration of the HO2 + NO reaction and the resultant increase in 

stratospheric OH abundances leads to changes in the HOx, NO., and CIx 

chemical cycles. In the lower stratosphere, the OH:He 2 concentration 

ratio is now determined largely by the reactions, 

OH + 0 3 - HO2 + 02 

HO2 + NO - OH + NO2 

with the net rate of ozone catalysis by HOx being limited by the reaction 

HO2 +03 -*OH +0 2 + 02 

In the upper stratosphere and mesosphere, these reactions are superseded 

by the atomic oxygen processes, 

OH+O-dH+0 2 

HO 2 + 0-* OH + 02 

Increased OH leads to more complete conversion of NO2 into HNO 3, 

reducing the stratospheric ratio of NO, to HNO 3. Hydroxyl radicals also 

generate ozone-active chlorine gases, Cl and CIO, from their stable counterpart, Ha; increased OH thus leads to a more important role for chlorine 

compounds in ozone-depletion calculations. Both the NO, and CLx interactive effects are discussed further in subsequent sections. 
40 .] 

35 

= 

'Nt EVANS el 81 (1978) 

o 

ACKERMAN et al. (1975) 

26 

:OSSERVE D 

20 

15 

0.111 

RATIO NO 2 1NO 

CALCULATED0 

10 

Figure 29. Comparison between observed and calculated [NO 2 ]/[N O ] ratio. 
ONE-DIMENSIONAL MODELS 157 

The larger HO2 + NO reaction rate results in an increased rate of ozone 

generation by smog chemistry. When HO, is produced during the oxidation of methane and carbon monoxide in air, it can react with NO to give 

an NO2 molecule that is rapidly photolyzed into NO and 0. The 0 subsequently reacts with 02 to form ozone. With the older, smaller HO2 + NO 

rate constant, HO2 was much more likely to react with 03 than NO, 

destroying ozone in the-process. OH concentrations in the troposphere are 

quite uncertain (see Chapter 3) and are of great significance in determining 

the efficiency of some of the more reactive halocarbons as sources of 

stratospheric Cl, . Tropospheric measurements and improved tropospheric

chemical modeling are needed to improve this state of knowledge. 

COMPARISON OF THEORY AND MEASUREMENT: NO x 

One-dimensional models have been shown to describe the broad-scale 

features of the observed distributions of the nitrogen-containing species 

in the stratosphere; however, a precise quantitative assessment of their 

reliability and accuracy is not available. One reason for this is the present 

lack of a sufficient number of accurate observations of the trace species. 

However, there is also the basic question of how the results of a 1-D model 

are to be interpreted. The models are intended to simulate average global 

conditions, but, because the number of observations is limited, it is not 

presently possible to select the proper specific set of observations for a 

global average comparison with the model. Furthermore, naturally occurring variations in the nitrogen oxides due to changes in season, latithlde, 

time of day, and stratospheric circulation will be reflected in the observations but are generally treated crudely or not at all in the model. 

The usefulness of 1-D models can be judged by their ability to satisfactorily describe the observed ratios of the nitrogen species and also the sum of 

all the nitrogen species. In this regard, the many observations of [NO] and 

[HNO 3 ] in the stratosphere below 25 kilometers are of only limited value 

for testing 1-D models because only one species density was usually measured at a time and because, of-the great natural variability expected in this 

region. For example, below 25 kilometers, large day-to-day and seasonal 

variations are expected in the concentration of ozone, HN03, and total 

odd nitrogen (NO + N02 + 2N2 0 + HNO3 + CIONO2 ). Accordingly, 

this subsection will concentrate on experiments in which two or more 

species were measured simultaneously in regions above 25 kilometers at 

which photochenical equilibrium is a good assumption for the nitrogen 
158 CHLOROFLUOROMETHANE WORKSHOP REPORT 

species. [NO 2] atd [NO] are in equilibrium with each other during the 

day with a time scale of approximately 100 seconds. Their ratio is given 

by 

[NO2] [03] " kCo + o3) 

[NO] JNO + [0] - k(4o2 +0) 

This ratio provides a good test for 1-D models;however, in addition to the 

difficulty involved in simultaneous measurements of [NO], [NO 2 ], and 

[031 , the theoretical ratio may be different from the observed one because 

of errors in the reaction rates or absorption cross section of NO, or because radiative scattering is neglected. Furthermore, NO2 is usually measured at sunset from its solar-absorption spectra so that partial conversion 

of NO into NO2 may occur and be reflected in the data. 

The most useful -measurement of this ratio to date is that of Evans et al. 

(1976) in which simultaneous measurements of [NO], [NO 2 ], [HNO3], and [O] were made from 10 to 35 kilometers on July 22, 1974, at 

58' N and the results compared with a diurnal integration of the photochemical equations. 

Figure 29 shows that the observed [NO2] I [NO] ratio of Evans et al. (1976) 

is roughly twice as large as they calculated with a diurnal model using the 

simultaneously measured ozone profile. They estimate the uncertainty in 

the observed ratio to be ±40 percent. A similar error is likely in the calculated ratio because of errors in the observed ozone concentration (±10 

percent), the reaction of NO and 03 (approximately 25 percent), and in 

JNO Although the agreement between the theoretical and observed 

[No2 ]/[NO] ratio is reasonable considering the uncertainties involved, this measurement must be interpreted with caution. 

As previously mentioned, the observations were made at sunset so that 

partial conversion of NO into NO 2 may have occurred. Furthermore, 

while NO is measured in situ, NO2 is measured along a tangent path and 

therefore may not be indicative of local conditions. In addition, the omission of albedo effects from the diurnal integration can have an important 

affect on the [NO, II [NO] ratio. 
ONE-DIMENSIONAL MODELS 159 

Also shown in figure 29 is the measured ratio of Ackerman et al. (1975) 

at approximately 300 latitude. The theoretical [NO 2 ]/ [NO] ratio cannot 

be computed for their data because ozone was not simultaneously measured. Because these are also sunset measurements, the same sources of 

uncertainty discussed with regard to the data of Evans et al. (1976) must 

be kept in mind. A range of calculated ratios from several models with 

self-consistently cdalculated ozone profiles are also shown in figure 29 even 

though they should not be directly comparable. These model results are 

all 24-hour averages. 

In conclusion, the attempts to measure the [NO2 ] / [NO] ratio to date are 

not yet conclusive tests of the theory, mainly because of the ambiguities 

associated with sunset measurements. A definitive test will be possible 

only when in situ NO 2 , NO, and 03 observations at times near midday are 

compared with diurnal integrations that include the effects of multiple 

radiative scattering. Actual solar-flux measurements, both direct and diffuse components, between 300 and 400 nanometers would also be quite 

valuable to compute JNO2 . 

It should be pointed out that the [NO 2 ]/[NO] ratio may depend on the 

CLx species if the CIO concentration is as large as that reported by Anderson (1971 and 1976). A small correction also exists for the recently 

measured fast rate constant for NO + HO2 -+ NO2 + OH of 8 X 1012 cm 3 

molecule' seC 1 (see Chapter 1). In this case, the ratio is given by 

[NO2 ] (O] .k(1 o + 03) + [C0] • k(co + NO) + [H021 " k(o1 2 + NO) 

[NO] JNO 2 + [01 - k(No2 +0) 

-3 If, for example, the CIO concentration is 5 X 108 cm at 30 kilometers,. 

the first two terms in the numerator would be of comparable magnitude, 

and the [NO 2] I[NO] ratio would be approximately 1.8 times its value 

when Cl. is not included. 

The ratio of [HN031 to [NO 2 ] also provides a useful test of 1-D models, 

although its theoretical value is less certain because it requires knowledge 

of the OH concentration. This ratio should be observed in regions of 
160 CHLOROFLUOROMETHANE WORKSHOP REPORT 

photochemical equilibrium (at low latitudes above 20 kilometers and at 

middle and high latitudes above 25 kilometers). The ratio is given by 

[IN03] k(o H +NO,) - [OH] * [M] 

[No2] JHNO3 + k(oH + HN0 3) [OH] 

It is important to note that the equilibrium between [N0 2 ] and [HNO] 

is established over a longer time scale than that between [NO] and [NO.] 

(approximately 1day at 25 kilometers versus 100 seconds), but this adjustment time is still shorter than characteristic transport times. The 

[HNO3]I [NO.] ratio observed by Evans et al. (1976) during the July 22 

flight and their calculated value obtained by diurnal integration for identical conditions are shown in figure 30. Also shown is the range of values 

obtained in the l-D models compared at the July update workshop. The 

ratio calculated by Evans et al. (1976) was in reasonable agreement with 

the measurements., However, because the OH concentration was calculated rather than measured, this agreement must be considered fortuitous. 

Since that time, updates of rate coefficients, especially NO + H0 2 , have 

significantly increased the calculated OH concentration leading to the 

HN0 3 /N0 2 ratios shown by the shaded band in figure 30. Although 

these are 24-hour averages, the substantial disagreement between the 

measured and calculated ratios is significantly larger than expected variations during a diurnal cycle. 

Observational validation of the [fNO,]/[NO2 ] ratio must await OH 

measurements in the 25- to 35-kilometer region. Nevertheless, simultaneous measurements of only [HNO3 ] and [NO2 ] are of value because they 

permit an indirect check of OH concentration derived from a model. 

Here again, [NO2 ] should be measured at times near midday and the 

effects of multiple radiative scattering included by calculation or direct 

measurement. 

The presence of N2 05 has not been established by direct observation 

although diurnal integrations in l-D models have shown that it should be 

an important species in the stratosphere. Recently, King et al. (1976) have 

attempted to establish an upper limit for the abundance of N2O5, but 

their results do not rule out the presence of N2 0 5 in substantial amounts. 
40 

35 

30 

25 

~EVANS at al. (1976) 

20 

0.01 01 1 

RATIO IHNO3 INO 21 

Figure 30. Comparison between observed and calculated [HNO 3 ]/[NO2] ratio. 

10 
162 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Table 29 lists the ratios [N2 0s]/[N0 2 ] and [N20 1 WNO 3] for times 

1 hour after Sunrise and 1 hour before sunset. These values were obtained 

with the model of Kurzeja (1975) except that they are for 300 latitude and 

include the effects of scattering on photodissociation rates. The ratios of 

[N2 0s] / [HNO 3] at sunset are much lower than the sunrise values and are 

within a factor of 2 of the upper limit ratios of IN2 0 5/ [HNO 3 ] quoted 

by King et al. (1976) for sunset conditions. 

Table 29 

Ratios of N2 0 to NO2 and HNO3 for Times I Hour 

After Sunrise (SR) and 1 Hour Before Sunset (SS) 

[N,0 5 j [NO] [N 2 051 [1r2o] N20 5 MLIng Ratios (ppbv) 

Kilometers [NOJ SR [NO2] SS [INO3] SR [HNO3 ] SS Sunnso Sunset 

20 033 0070 007 002 035 0.10 

22 04S 0.088 0.13 0.93 0.65 0.15 

26 05 0.088 0.4 0-08 IAO 028 

30 078 0.052 0.8 0.1 02 0.25 

34 032 0.14 IA 005 1.4 0.05 

The concentration of N2 05 can also be inferred from Noxon's (1976) 

measurements of the diurnal variation of the NO 2 column content. Because NO2 is converted into N205 at night, the NO2 decrease (when 

divided by 2) should be equal to the production of N20,. Analysis of 

Noxon's results indicate an N205 mixing ratio of approximately 1 part per 

billion by volume (ppbv), which is consistent with diurnal models (table 29). 

Thus, N20, mixing ratios of approximately 1 ppb that are indirectly supported by Noxon's NO2 observations cannot be ruled out by the upper 

limits of King et al. (1976). Because N2 0 5 acts as a significant reservoir 

for NO., efforts to determine its concentration shoud continue. 

Although the ratios of the nitrogen oxides are maintained by local photochemical conditions, their sum, which is defined as odd N (odd N = NO + 

NO2 + 2N2 05 + HNO3), is highly transport-dependent. Odd N, produced 

mainly by the reaction N2 0 + 0(D) -.-2NO in the stratosphere between 

20 and 45 kilometers, is transported downward into the troposphere and 

upward into the mesosphere. The amount of odd N is an important 

quantity in the CFM problem for two reasons: (1) large NO2 concentrations increase the rate of ClONG 2 production, and (2) large NO concentrations decrease the CIO to Cl ratio. Both effects work to decrease the calculated ozone-colunm reduction. Small odd-N concentrations will have the 

opposite effect. 
ONE-DIMENSIONAL MODELS 163 

The odd-N distribution depends on the eddy-diffusion coefficient that 

governs the rate of N2O input into the stratosphere and the removal rate 

of odd N from the stratosphere. It also depends on the photodissociation 

rate of N2O, the concentration of 0(D), and the rate of the reaction N20 

+ O(QD) ->2NO. The odd-N concentration in the upper stratosphere depends on the dissociation rate of NO. The uncertainty associated with this 

calculation is illustrated by comparing the models of Crutzen and Isaksen 

(preprint, 1976) and McConnell and McElroy (1973). The N20 distributions, odd-N production rates, and odd-N distributions are shown in figures 

31, 32, and 33, respectively. As can be seen from these figures, Crutzen 

and Isaksen's larger odd-N concentrations result from larger N,0 mixing 

ratios. Their O( 1D) concentrations are also roughly 20 percent larger than 

those of McConnell and McElroy. The difference between the two N20 

distributions is due to Crutzen and Isaksen's larger eddy-diffusion coefficient below 30 kilometers and to their use of smaller absorption cross 

sections for N.O. Although Crutzen and Isaksen's N2 0 profile is in 

better agreement with observations below 35 kilometers, neither profile 

can presently be ruled out by observations. 

In figure 33, the observed odd-N mixing ratios of Evans et al. (1976) and 

Ackerman et al. (1975) are shown. The Ackerman profile was obtained by 

adding their measured NO and NO2 amounts to the calculated HNO3 

mixing ratio of Crutzen and Isaksen (1976). Although neither includes 

N2O5 , this neglect will not seriously affect the results because the measurements were made at sunset when most N2 0 5 has been converted into NO 

and NO2 . The observations suggest an odd-N mixing ratio of 1.1 to 

1.4 X 10-8 above 30 kilometers, which is consistent with the predictions of 

most 1-D models. Also shown in figure 33 is a range of results for total 

odd N above 35 kilometers obtained by eight different models at the July 

1977 update workshop. 

In the lower stratosphere, comparison between theory and observation is 

more difficult because of the large natural variability in odd N that accompanies the steep gradient in mixing ratio. Comparison between theory and 

observation must await a greater number of measurements to establish 

time and latitudinal averages. 

Other sources of odd N have been suggested. Figure 32 shows the source 

due to cosmic rays as derived by Nicolet (1976) for latitudes greater than 

600 and the source due to solar-proton events (Crutzen et al, 1975). The 

solar-proton event (SPE) of 1972 (the largest described by Crutzen et al.) 
N20 

40 0 

0 

40 c 

N0 

100 

10.7 10' 10-8s 

MIXING RATIO (volume) 

Figure 31. Comparison between N20 distributions as calculated by McConnell and 

McElroy (1973) (M) and Crutzen and Isaksen (preprint, 1976) (C). 
40 

SOLAR-PROTON 

EVENT 

M C 3 

S30 

20 -

0 

N 

100 101 

ODD-N PRODUCTION (cm- 3 sec - 1) 

102 103 

Figure 32. Comparison of odd-N production rates of McConnell and McElroy 

(1973) (M)and Crutzen and Isaksen (preprint, 1976) (C). 
CC 

40g 

EVANS et ai le ACKERMAN et al. (1975) 

30 -­

" 10 -08 10 7 

MIXING RATIO (volume) 

Figure 33. 0bserved odd-N mix!ng ratios of Evans et al 11976) and Ackerman et al. (1975). 
ONE-DIMENSIONAL MODELS 167 

has been plotted as a yearly average NO production rate under the assumption that there were no other significant solar-proton events in 1972. The 

SPE production rates have also been divided by 6 to convert them to a 

hemispheric average value. 

The relatively small NO production from cosmic rays and the low level of 

input suggests that this source will have very little effect on odd N in the 

middle and upper stratosphere where ozone destruction is most important. 

Similarly, NO production from solar-proton events will not have a significant effect on the long-term odd-N budget, but may have an important 

short-term effect. These two sources of odd N are not included in 1 -D 

models designed for studying the CFM problem, and, considering the small 

magnitude of these sources compared with production from N2 0, this 

neglect seems reasonable. 

One-dimensional models do not include tropsopheric sources for stratospheric odd N. This is based on the assumption that washout and rainout 

are rapid enough to prevent upward fluxes into the stratosphere and also 

on the observed increase in the HNO3 mixing ratio in the lower stratosphere with height that suggests a downward flux of odd N. Tropospheric 

observations ofodd-N species are necessary to confirm this assumption. 

Most 1-D models yield maximum odd-N mixing ratios of between 12 and 

19 ppbv, which is consistent with the few observations of odd N presently 

available. However, there is much room for improvement, particularly in 

predicting the magnitude of the maximum odd-N mixing ratio and in 

verifying it observationally. More frequent observations of NO, NO2, and 

N20 above 30 kilometers would be a great aid in establishing the average 

odd-N mixing ratio and its variability with time. 

Modelers should strive for more careful calculation of the dissociation 

coefficient of NO, which is important for odd-N loss above approximately 

35 kilometers. They should be aware that the frequently used JNO values 

of Cieslik and Nicolet (1973) need correction. (The dissociation rate for 

their 6 (1-0) band should be reduced by a factor of 1/2.2 to make them 

consistent with the band strengths of Bethke rather than those of Caflear 

and Pilling (see Cieslik and Nicolet, 1973)). 

The observed N70 mixing ratio in the troposphere is used as the lower 

boundary condition for 1-) models, and additional measurements would 

help determine which value is most representative in the 028- to 0.33­

ppmv range currently used by modelers. 
168 CHLOROELUOROMETHANE WORKSHOP REPORT 

This discussion has concentrated on those areas in which 1-D models-can 

be tested with observations so as to minimize ambiguity. In regions in 

which transport is very important and the 1-D diffusion assumption is 

most questionable (i.e., the stratosphere below approximately 25 kilometers), not much is to be gained by demanding good a prioriagreement 

between the observed and predicted nitrogen-oxide concentrations. Effort 

is better spent ensuring that 1-D models agree with observations in the 

region above 25 kilometers in which transport is less important. 

Major efforts should be devoted to deciding whether N.0 is the major 

source of odd N and to establishing the size and variability of odd N in the 

region above 30 kilometers. This will require more extensive measurements of [N 2 0] , [NO], and [NO 2] at 30 kilometers and above. 

Carefully planned measurement programs of the densities of several species 

simultaneously, such as that of Evans et al. (1976), should be encouraged. 

However, significant improvement in deciding the ability of 1-D models to 

predict the ratios of the NO2 species cannot be expected until techniques 

are devised for in situ measurement of [NO2 ] at times other than sunrise 

and sunset. Efforts to observe N2 0 5 should continue because diurnal 

integrations invariably indicate its importance as a reservoir for NO and 

NO2 ' 

COMPARISON OF THEORY AND MEASUREMENT: CIO, 

INTRODUCTION 

A useful distinction is that between organic and inorganic molecules containing Cl atoms. As far as is known, the major sources of C1 atoms for the 

stratosphere are "organic'* compounds (e.g., CF2C12 , CFCl3 , CC14 ,CH 3C1, 

and CH 3CC13) originating at ground level; one exception is the future 

Space Shuttle HCI injections. Sources are discussed in the following 

section. These organic molecules do not affect stratospheric [0] directly, 

but only through their photo-oxidation products, which are inorganic molecules containing Cl atoms (e.g., Cl, CIO,HC1, and CIONO2). The term 

"odd chlorine" has been coined by aeronomers, in analogy with odd nitrogen, odd oxygen, and odd hydrogen; odd chlorine is equivalent to "inorganic chlorine" in this discussion. The notation "C1X" is widely used to 

signify odd chlorine, whereas the term "ClO" is used to lump together Cl, 

C1O, and any higher oxides. It should be noted that the C1X grouping can 

contain any and all Cl-containing molecules that are produced in the 
ONE-DIMENSIONAL MODELS 169 

stratosphere from the photo-oxidation of organic chlorine and subsequent 

reaction, i.e., the concentration, [CIX], of odd chlorine = [ClJ + [C1OJ + 

[HI] + (CIONO2 ] + [any other inorganic Cl-containing species]. Finally, 

the mixing ratio (by volume) of any species is simply its molefraction. 

SOURCES OF STRATOSPHERIC ODD CHLORINE 

Since Molina and Rowland's first papers (1974 and 1975) on the stratosphere as a sink for the chlorofluoromethanes, CF2C12 , CFCI3 , and CCI 4 , 

a great deal of thought has been given to other compounds, natural or anthropogenic, that can furnish Cl atoms to the stratosphere, It now seems 

clear that CH3 C1 (probably natural) furnishes 40,000 to 250,000 metric 

tonsofCl to the stratosphere per year. This figure is based on Rasmussen's 

(1976) data, Lovelock's (1976) data, recent measurements by Singh et al. 

(1976), and from calculations. By comparison, the yearly stratospheric input of Cl atoms due to 1975 usage of CF2 Cl. and CFCl 3 (with no tropospheric losses) will be about 500,000 metric tons. As a metal cleaner and dogreaser, methyl chloroform (CH 3CCl3 ), when lost to the atmosphere, 

probably now represents a 30,000-metric-ton-per-year source (Crutzen and 

Isaksen, 1976), and its usage is growing. Contributions from manmade 

C2C14' C2HC3' CH 2Cl 2 , and CHCI3 also may not be neglected, as approximately 0.5 to 3 percent of ground-level emissions can enter the stratosphere bearing Cl atoms. In calculating the flux of each organic Cl form 

that can enter the stratosphere, the tropospheric decomposition rate is 

found to be uncertain to some extent. For CF 2C12 , CFC13 , and CCl4 , the 

NAS/National Research Council (NRC) report discussed unknown but 

possible tropospheric removal processes. For the other organic Cl molecules (or chlorocarbons), attack by OH in the troposphere (usually Habstraction) is an important tropsopheric sink. The great uncertainty in 

tropospheric OH concentrations (easily 1 power of 10 even in models that 

neglect variations in relative humidity, tropospheric 03, and incident ultraviolet) dictates that, when attempting to estimate stratospheric Cl inputs 

from, for example, CH 3 C1 and CH3CC13 , the estimate should be parametrically represented with a range of tropsopheric [OH] . The possibility 

of ground-level inorganic sources influencing the stratosphere does not 

appear to be large, but the tropospheric behavior of HCI on which this 

dismissal is based has not been firmly established. (See the discussion of 

HCi tropospheric data in the following section.) 

These known sources of stratospheric CIX can be used in atmospheric 

tracer models to predict C1X concentrations. To date, only 1-D models 

have been employed toward this end. In these studies, it is assumed that 
170 CHLOROFLUOROMETHANE WORKSHOP REPORT 

atmospheric CH3 Cl concentrations are in steady state with CH 3Cl sources 

and sinks. Time-dependent calculations are required for the CF 2 C12 and 

CFC13 inputs because of their history of increasing usage. In such timedependent calculations, the vertical eddy-mixing coefficient controls the 

rate of upward mixing from ground level through the tropopause and to 

the levels above 25 kilometers at which most of the CF 2 C12 and CFC 3 is 

photolyzed. It is therefore obvious that the stratospheric CIX profile predicted for a given year will vary from model to model, not only in CIX 

peak concentration but also in profile shape. The range of CIX predictions 

for late 1976 from several scientific groups is from 1.3 to 2.7 ppbv. Figure 

34 shows a calculation for a mid-1976 stratosphere with an eddy coefficient twice as large as that of Hunten (1975). Contributions to the calculated CIX profile from CF CI, 21 CFCI3, CC14 , and CH 3Cl are noted. In 

this calculation, the steady-state assumption was applied to CC14 , and it 

was assumed that all four Cl atoms from each CC14 are released (i.e., that 

CCl2 0, etc., are quickly photo-oxidized). The peak mole fraction, 1.9 

ppb, is attained by about 35 kilometers. Its value would be higher, 

between 2.0 and approximately 2.5 ppb, if CH 3 CCI3 and other chiorocarbons were included. It would be lowered if CC120 were assumed to be 

more stable, if a value less than 0.95 ppb were taken for a global average 

CH3CI tropospheric background, and if a slower eddy-mixing rate were 

assumed. 

It should be reemphasized that the CIX concentration, by definition, is the 

total concentration of inorganic molecules containing Cl. Thus, if and 

when a stratospheric concentration of CIO or HCI, for example, is measured to be greater than about 2.5 ppb, that measurement cannot have been 

predicted using 1-D models (based on seasonally averaged tracer data, at 

best) and the C1X sources just discussed. The highest CIO concentrations 

in the range Anderson (1971 and 1976) has observed are-perhaps three 

times higher than simple l-D calculations can yield with known CIX 

sources. Three obvious possible explanations are: (1) that meteorological 

variability leads to factors of 3 or more swings in the C1X concentrations 

above 25 kilometers, (2) that as yet unidentified CIX sources are significant, and (3) that Anderson et al. (1977)have measured CI from contamination (in their instruments or from, for example, solid-fuel rocket plumes 

not associated with their program) or have detected nonresonant (with 

C1) photons. 

Possibility (3) above has been investigated during the design, fabrication, 

testing, and flight of the Anderson et al. (1977) instruments and seems 

very unlikely. Indeed, no plausible criticisms currently exist. Possibility 
60 

I I 

C 2 CI 

I I 

INPUTS FOM 

CFCI3 04 

I I I 

H3

C I TOT L. 

I I I I I II 

60 

s4O 

f 

30 

I 

Iz 

z 

7 

20 I 

af 

I--"" /zi. 

100 

loo 

10-10 

I I I Ii I I il 

10 

MOLE FRACTION 

INORGANIC ODD CILORINE 

I I I I lI iI 

108 

Figure 34. Calculated CIX mode fractions using a mid-1975 stratosphere with a 

vertical transport coefficient twice as large as Hunten's (1975). -
172 CEILOROFLUOROMETHANE WORKSHOP REPORT 

(2) is being explored by many investigators; it will no doubt receive even 

more attention if, as the CIX data base grows, there appears to be a convergence of CIX concentrations to mean values above, for example, 3.5 

ppb. Possibility (1) will require the attention of stratospheric dynamicists for quantitative theoretical assessments, but measurements of the various CIX species at different seasons and latitudes-will permit meaningful 

averaging of the variability. Seasonal, midlatitude averages are probably 

best for comparison with 1-D model predictions. 

MEASUREMENTS AND THEORETICAL PREDICTIONS FOR 

CIX SPECI ES 

HCI
 

Because HCI in the stratosphere should have a photochemical lifetime of 

hours, its expected variability is less than that for Cl or CIO (photochemical lifetimes of seconds), and, thus, it may be a candidate for comparisons 

between measured and theoretically predicted concentrations. Figure 35 

presents such a comparison. The experimental data are the same as those 

given in figure 10 with the addition of the recent experimental profile of 

Eyre and Roscoe (1977), labeled ER in the figure. Several of these profiles indicate a maximum in the HC1 concentration between 25 and 35 

kilometers with a possible turnover above. 

Sample theoretical curves for HCl mixing ratio with altitude, employing 

the latest rate information for the reaction of HO2 with NO (see Chapter 

1) are shown in figure 35 for two asymptotic values of CIX, 1.6 and 3.2 

ppbv. These show a slight flattening of the distribution between 20 and 

35 kilometers but no maximum. 

The variation between different measurements is comparable to the differences between measurement and observation. Thus, it appears that the 

budget of stratospheric CIX may be reasonably represented by the view 

previously outlined. (See figure 34.) On the other hand, it is possible to 

be very critical of the HC1 measurements to date. For example, the filter 

collections of Lazrus and coworkers actually measure any substance 

(probably acidic) that yields C1" in aqueous solution. Also, Williams et al. 

1 (1976) have noted that the HC1 absorption at 2925.9 cm" (the line used 

1 in all reports thus far) and its 37 C1 counterpart at 2923.7 cm- appear to 

be blended, thus implying that the [HCI] deduced from these absorption 

lines may be higher than is actually present. Further, the R5 H37 CI line " at 2995.8 cm 1 that should be observed (it lies in a clear interval) has not 

been seen. 
ONE-DIMENSIONAL MODELS 173 

IoI i" J I I" I I I i I I I i I I 

40I 

30t 

20
 

. CALCULATION 

16 pplv CI. 

and 3.2 ppbv CI 

to 

0 05 10 1i 20 

MOLE FRACTION (109) 

Figure 35. Comparison between measured and 

prediction HCI concentrations. 

Nevertheless, the infrared data do appear to identify stratospheric HC1 

and, together with the filter paper data, do show that its mole fraction increases, with altitude up to about 30 dlometers, consistent with a middle 

stratospheric source of CIX such as the photo-oxidation of CF 2Cl2 , CFC13, 

CH 3Cl, and CC14. Measurements to higher altitudes (for example, 40 to 

45 kilometers) would be very helpful in deducing the total C1X in the stratosphere. 

Some mention should also be made of the tropospheric data and predictions for HC1. The view that, in the troposphere, HCrs mole fraction decreases rapidly because of rainout effects is fairly widely accepted now, 

but tho evidence for this view is not as strong as might be desired. The HC1 

tropospheric scale height deduced by Fanner et al. (1976) ("F" in figure 

35) is approximately that predicted from the "H 20-control" argument 

that follows, and it leads to the very small HC1 mole fractions (< 10"

10) 

predicted and measured just above the troposphere. On the other hand, 
174 CHLOROFLUOROMETHANE WORKSHOP REPORT 

the feature referred to as the RI line of the H3 7 C1 fundamental by 

Farmer et al. (1976) appears as a doublet in the recent spectra of Bradford 

et al. (1976), neither line of which is positioned properly for HCI (A. Goldman, private communication, 1976). Thus, there is a need for corroborative studies. Similarly, the early analytical procedures of Junge (1957), 

Duce et al. (1965), and others may have permitted particle-to-gas conversion in the inlet systems of the apparatus, a criticism offered by Stedman 

et al. (1975) and by Berg and Winchester (1976). In that case, the groundlevel HC1 mole fraction of I to 2 ppb deduced by these workers would be 

an upper limit. Fortunately, the ground-level concentration of HC1 has virtually no effect on stratospheric calculations if rainout is indeed a rapid 

sink from HC] in the lower troposphere. The actual rainout rates are not 

well-known, but the available data appear to be consistent with a removal 

rate fast enough to destroy any coupling between surface HCI and the 

stratosphere. 

The suggestions and evidence just cited have been used to dismiss the influence of ground-level sources of CIX on the stratosphere (e.g., Cicerone 

et al. (1975); Cicerone (1975); and NAS (1976)). In current models, 

CIX is removed in the troposphere by assuming a heterogeneous loss term 

with - a rate of approximately 1.7 X 106 sec-1 (ie., atime constant of the 

order of days). The picture that emerges fits available data on HCI, but, 

as has been noted, neither the available data nor the theory is compelling. 

Thus, there will be some continuing uncertainty as to whether ground-level 

or tropospheric CIX or NOx sources exert any control over stratospheric 

C1X and NOx budgets. 

CI and C10 

The recent measurements by Anderson et al. (1977) by a spectroscopically 

specific technique (Cl atom resonance fluorescence and chemical conversion of CIO) are the only published or about-to-be-published data. Their 

results are discussed in Chapter 3 of this report, as is the comparison with 

noontime CIO and Cl predictions prior to the revised HO2 + NO reaction 

rate. As previously mentioned, the C10 data from two flights indicate perhaps 5 ppb CIX in the stratosphere assuming model calculated ratios for 

[CIO]/ [ClX]. If there were at least 5 ppb CIX, then variations in [OH] 

(Burnett, 1976) would cause variations in the [CIO] I [HCI] ratio and the 

observed [C1o]. On the other hand, 5 ppb of CIX cannot be accounted 

for from known sources, and the agreement between HC1 predictions and 

measurements would be stretched so that the uncertainty limits on the 
ONE-DIMENSIONAL MODELS 175 

HQl measurements and theory would no longer overlap. All of the possibilities discussed in the section, "Sources of Stratospheric Odd Chlorine" 

must be explored, with care taken to calculate diurnal variations in Cl and 

C10 and to employ proper seasonal and latitudinal groupings. Measured 

ratios of [CL]I[CIO] can also be compared with theory, but most of the 

observations will be dependent on the -variability in other species; e-g., 

[0] [0. 1. Above approximately 35 kilometers, the 0103 ratio determines 

the CICIO ratio. 

A comparison of current Cl and C10 measurements with theory is shown 

in figure 36 for several model calculations. The revised rate constant for 

HO2 + NO of 8 X 10-12 cm 3 molecule sec 1 is included. This has the 

effect of increasing the amount of C10 below 35 kilometers leading to 

better agreement with the intermediate experimental profile in this region. 

In all cases, the experimental CIO falls off more rapidly with altitude above 

35 kilometers than is theoretically predicted. CIO at 35 kilometers increases by between 12 and 50 percent for a range of models relative to the 

case with the earlier HO2 + NO rate constant. 

s0
 

45 

40K 

D A

25 

20
 

20 -A 

is I I 

105 10.6 108 150 109 

CONCENTRATION (omf 3 ) 

Figure 36. Comparison of calculated and measured CO and Cl concentrations. Measurements are those of Anderson and coworkers (see figure 12]. 

Calculations are: A - Ames/RDA model with 3.2 ppbv CIX; D - Dupont 

model with 1.9 ppbv CIX. 
176 CHLOROFLUOROMETHANE WORKSHOP REPORT 

CION02 

Murcray et al. (1977) have reported upper limits to the CIONO2 abundance in the stratosphere of 0.8 ppb at 25 kilometers increasing to 2 ppb 

at 30 kilometers. Further, these numbers were based on absorption coefficients that according to the authors may be a factor of 2 too low. If they 

are, then the upper limits cited would be reduced proportionally. Comparisons can be made with predictions fox September 1975 when there 

should have been between 1 and 2 ppbv CIX. Doing so leads to calculated 

concentrations approximately eqtaLrto or slightly less than the upper 

limits of Mureray et al. (1977). If these upper limits are a factor of 2 

lower, then the calculations exceed the upper limit at 25 kilometers and 

approximately equal the upper limit at 30 kilometers. 

Theoretical calculations of ClONO2 concentrations in the stratosphere 

currently include destruction by sunlight and by reaction with 0 atoms, 

with products assumed to be one CIX and one NOX molecule. Attack by 

OH appears to be noncompetitive, and heterogeneous decomposition, 

although seen on some laboratory surfaces, is neglected, Its rate of formation according to reaction 

k CIO + NO 2 + M * C1ONO2 +M (R1) 

obviously depends on [ClO], [NO 2 ], and k. Thus, there are possible CIOpoor or NO2 -poor altitudes, times of day, or future C1X-concentration 

scenarios wherein the rate of formation is limited by the CIO or NO2 

available. Certain of these can spawn nonlinear effects in the calculated 

effect of ClONO2 on stratospheric chemistry as, for example, CIX or NO, 

is added to the stratosphere. 

The lack of laboratory measurements of the pressure dependence of k 

above 5 torr introduces another uncertainty into ClONO2 calculations. 

This is serious because the limited data for atmospheric CON0 2 and the 

sensitivity of calculated CIONO 2 levels to the method of calculation already combine to represent a substantial uncertainty in 1-D model calculations of fluorocarbon effects. There is a-clear need for higher pressure 

determinations of k. Another important question is whether reaction RI 

can have alternate reaction paths not leading to ClONO2 , a possibility not 

yet ruled out. Similarly, it remains necessary to determine the products 

of ClON0 2 photolysis. 
ONE-DIMENSIONAL MODELS 177 

Summary of CIX Measurement and Theory 

Several things have been learned from the three sets of flight data (Anderson et al., 1977) on stratospheric CI and C10. First, the detection and 

measurement of CI and CIO, coupled with previous measurements of 

0(3P) and with laboratory kinetic studies of the reactions 

C1 + 03 -+C + 02 

CIO + 0 C1 + 02 

establishes that catalytic removal of 03 by these reactions is proceeding in 

the high stratosphere. The measurements to date do not establish the C10 

catalytic chain length because of lack of knowledge of total CIX concentrations in the present stratosphere and because of the complications in 

the middle stratosphere involving ClONO 2 . 

Also, as mentioned earlier in this chapter and in Chapter 3, two of the 

flights of Anderson et al. (1977) (July and October 1976) found C10 concentration profiles whose peak values at local noon equaled or exceeded 

the total CIX concentrations predicted from known sources of CIX Thus, 

even 100 percent conversion of HC1 to C10 under noontime conditions 

around 35 kilometers cannot furnish CIO concentrations as high as the 

midrange of the Anderson et al. measurements. Only the October 1976 

flight showed peak C10 mole fractions of about 3 ppb CIX. The possibility must be considered that there is, for example, 5 ppb CIX in the middle-to-upper stratosphere and that the measured variability of [C10] so far 

is due to variations of [OH] and [HO2 ] (see discussions on HOx), representing variations in the ratio [CIO]/ [HCI] at high altitudes and [C10]/ 

([C-1] + [ClN0 3]) at lower altitudes. The consequences of thispossibility 

must be examined (i.e., that there is now a CIX mole fraction of about 5 

ppbv CLX). It would be necessary to consider immediately the possible 

origin of 5 ppbv CIX, more than half of which is not accounted for with 

known sources (see the section on "Sources of Stratospheric Odd Chlorine"). This hypothetical excess of "-3ppbv ClX could be due to natural 

sources not yet identified. If so, those sources must be identified to understand all relevant physics and chemistry. It may be possible to predict 

the ozone perturbations due to CFM usage simply by evaluating the 0 3 

changes, starting from a stratosphere with 2 to 5 ppbv C1X. However, this 

would be difficult in the absence of knowledge of the total stratospheric 

chlorine budget and distribution. If, on the other hand, the hypothetical 

excess of -3 ppbv CX in the 1976 stratosphere is due to anthropogenic 
176 CHLOROFLUOROMETHANE WORKSHOP REPORT 

emissions of some Cl-containing molecule, organic or inorganic, then the 

following questions must be answered: When did these emissions occur? 

Should they have led to noticeable 03 perturbations? What molecules and 

processes are involved in the emissions? When identified and studied, 

would these additional emissions lead to significant improvements in the 

understanding of atmospheric chemistry and physics? Whatever the origin 

of the hypothetical excess ClX, it would also lead to less calculated CH4 

near 40 kilometers than is measured. 

Finally, the first measurements of stratospheric C1 and ClO have led to a 

strengthened realization that measurements oftotal C1X in the high stratosphere axe needed. As noted previously, simultaneous measurements of 

species that control ratios such as C1ICI0 are also clearly required. More 

data for C10 and HC1 above 35 kilometers are important, as well as an 

explanation of why both, species appear to decrease more rapidly with 

altitudes above 35 kilometers than is theoretically predicted. 

TROPOSPHERIC AND STRATOSPHERIC SINKS FOR 

CHLOROFLUOROMETHANES 

INTRODUCTION 

A large body of information on the effects of chlorofluoromethanes, 

CFC!3 (F-l) and CF 2Cl 2 (F-12), on stratospheric ozone has been assembled and assessed by the Panel on Atmospheric Chemistry, National Academy of Sciences, Washington, D.C. The panel's findings are detailed in the 

recently released report, Halocarbons: Effects on Statospheric Ozone 

(NAS, 1976). A variety of tropospheric and stratospheric removal processes for F-I 1 and F-I2 were investigated by the panel. No significant sinks, 

other than the active removal in the stratosphere originally proposed by Rowland and Molina (1975) and Molina and Rowland (1974) have been 

identified. (Active removal processes deplete stratospheric ozone; inactive 

ones do not.) The conclusion of the panel is summarized in table 30. 

Among the inactive removal processes considered, only dissolution in the 

oceans has lifetimes (lower limits of 70 years and 200 years for F-1I and 

F-12, respectively) short enough that its maximum effect would be significant (a 20-pexcent reduction of predicted ozone destruction). A lower 

limit of 102 years for reactions with neutral molecules in the troposphere 

was based on the detection limits of the reaction rates with OH radicals in 

laboratory experiments and is likely to be much too short. Material 

balance has been made for F-11 and F-12 using estimated atmospheric 
ONE-DIMENSIONAL MODELS 179 

Table 30
 

Removal Times for F-Il and F-12
 

Process Removal Time (r) 

ars) 

Active removal in stratosphere 

Photolysis and 0('D) reaction 50;90 

Surface processes 

Removal by oceans > (70,200) 

Removal by soil and microbes > I0 

Entrapment in polar ice > 10 

Tropospheric processes 

Thotodissocwtion > 5X 103 

Reactions with neutral molecles >> 100 

Direct ionization ): 106 

lon.molecule reactions > 10 

Heterogeneous processes > 6X 104 

Lightning > 106 

Thermal decompositon > 104 

Inactive removal in stratosphere 

Ionic processes > I0 

Heterogeneous processes > zo, 

release and actual atmospheric measurements. The results can be interpreted, with some uncertainty, as consistent with little or no inactive 

removal for F-I 1. 

Recent laboratory measurements (Ausloos et aL., 1977) of the photolysis 

of F-1 I adsorbed on sand indicate the possibility of tropospheric removal 

when contact is made with sand. Quanlitative evaluation is not yet possible. 

In the following subsections, some of the removal processes listed in table 

30 (taken from the NAS report) and other relevant information not included in the panel report are discussed. 
180 CHLLOROFLUOROMETHANB WORKSHOP REPORT 

REMOVA&L BY OCEANS 

The flux from the atmosphere into the oceans for F-11 and F-12 was estimated with a diffusion model formulated by Liss and Slater (1974): 

D
 

F = - AC,
 

where D is molecular diffusivity, Z is film thickness, and AC, = Ceq - Cw 

is the difference between equilibrium and actual concentration in ocean 

water. = The flux would be maximum when AC1 ,

Ceq - aC0 where q is 

the solubility and C. is the atmospheric concentration. The removal time 

is then 

N N 

6Fmax eD aC o 

z 

where N is column density and e is fractional surface of oceans on Earth. 

The panel obtained r > 70 years for F-1i and r > 200 years for F-12. 

However, lunge (1976) calculated r > 208 years for F-11. The discrepancy resulted from the numerical values used for D, Z, and a. Table 

31 compares the two sets of numerical parameters. 

Table 31 

and F-2 a Oceanic Removal Times for F-1 

D z nt IT 

(i0 5 CmI $-') (10-3 cm) (t01$ molecules cmf " atmic ) yea Note 

12 4 4A(F-11),1.6(F-12) 70(F-I1),>200(F-2) b 

1.0 7 2.5 (F-11), 1(-12) 270 (F-11),> 600 (F-12) 

Calulad fat N =2X l0 moteculesft cm C --= 10 9atm.e-2/3,andat 288 K. 

bNAs (1976). 

CJung (1976). 

Rasmussen et al. (1976) recently reported an average concentration of 

0.17 X 10-9 grams Iiter "1 for F-11 in s'utface.ocean water under an atmospheric mixing ratio of approximately 120 parts per trillion. This concentration is higher than the equilibrium concentration using either Zeininger's 

solubility measurements (Junge et al., 1976) (by a factor of 2 to 3) or 

Lovelocles (by a factor of 1.5): 
ONE-DIMENSIONAL MODELS 181 

It is clear that more studies and measurements are needed to obtain better 

estimates of the oceanic removal times for F-il and F-12. 

PHOTODISSOCIATION IN THE TROPOSPHERE 

Rebbert and Ausloos (1976) recently reported a photodissociation cross 

section (or absorption cross section if quantum yield of photodissociation - is unity) of 3.7 ± 0.4 X 1026 cm 2 at 313 nanometers for CC14 at 300 K. 

Furthermore, the absorption cross section of CC14 on the long wavelength 

edge of the near ultraviolet continuum shows exponential dependence with 

wavelength. Exponential extrapolation fits Rebbert and Ausloos' cross 

section at 313 nanometers very well. Because the shape of the absorption 

bands for CC14 , CC13F, and CC12 F2 are similar, with a red shift from 

CC12F2 to CcI3F to CC14, exponential extrapolation is likely to give reasonable estimates of cross sections for CC13 F and CCI2F2 as well. Cross 

sections of the order of 10-25 cm2 at 280 nanometers and 10-28 cm2 at 

313 nanometers for CC13F are obtained by extrapolating the cross sections measured by Robbins and Stolarski (1976) and by Chou et al. (in 

press), corresponding to a removal time of the order of 104 years. The 

cross sections for CCl F2 in the 300-nanometer region are about two 

orders of magnitude smaller than those of CCI F. 

REACTION WITH OH RADICALS IN TROPOSPHERE 

Reaction R2 is endothermic by about 2 kcal-mol 1 (Foon and Tait, 1972; 

and Benson, 1968): 

CCl3F + OH -- CCl 2F + HOC1 (R2) 

Therefore, the reaction is expected to be too slow at tropospheric temperatures to be a significant removal process (NAS, 1976). However, OH 

radicals could be formed with vibrational excitation in reaction R3, which 

is the major tropospheric source of OH radicals. Reaction R3, yielding OH 

(v = 2), 

('D)+H 2 0- OH+ OH (v--2) (R3) 

is exothermic by approximately 8 kcal-mol "1 and could be a fast reaction. 

Rate constants for OH (v > 0) with CC3 F or CC2 F2 have not been reported, but the concentration of OH (v = 2) can be estimated by a simple 
182 CHLOROFLUOROMETHANE WORKSHOP REPORT 

kinetic scheme. Assume that all OH radicals in the troposphere are formed 

initially at v 2, followed by deexcitation and eventual removal by chemical reactions, 

kq OH (v=2) +M 4 OH+M (R4) 

kr 

OH + R - products (R5) 

Important quenchers (M),reactants (R), and relative rates, (k [M] ir [R]), 

are listed in table 32. The steady-state concentration of 6H(v = 2) is 

k [R] 

[OH(v = 2)] = kq IM] [OH]

Tropospheric concentrations of 106 - 107 molecules cm 3 have been measured for OH (Wang et al., 1975; Davis et al., 1976; and Perner et a, 1976). 

The upper limit for the concentration of OH (v = 2) is then 

3 [OH (v = 2)] = -12 3 X 107 = 1 molecule mcm

10-1


This concentration is too low to contribute significantly to the inactive 

removal of F-1 and F-12. 

REMOVAL BY PYROLYSIS 

There is evidence that F-11 in air can be pyrolyzed easily on hot surfaces 

(Smith et al., unpublished). All processes that use large quantities of air at 

high temperatures, such as combustion and. burning, need to be examined 

for possible contribution to the inactive removal of F-1I and F-12. The 

rate of removal for F-1I and F-1 2 by motor vehicles can be estimated reasonably well. For 1976, 1.1 X 1012 vehicle miles traveled in the United 

States was projected by the Department of Transportation (Clewell and 

Koehl, 1974). At an estimated average speed of 40 miles/hour' , airflow 

rate into a typical (6-cylinder) engine is approximately 1 X I os £ year-' . 

Gasoline accounts for approximately 20 percent of the fuel (oil, gas, and 

coal) consumed in the United States. The worldwide consumption of fuel 

may be twice that of the United States. 
ONE-DIMENSIONAL MODELS 183 

Table 32
 

Quenching of OH (v-2) and Reactions of OH
 

Nk 

molecules cm-3 s' 

q 

References 

[M] k [M] 

N2 andO0 <10 -13 b c <-10- 13 

H20 j1Qd e 10 " 2 10 "11 

R k 

V ____ [R]3 kc[R] 

molecules cm 3 s " 1 References 

CH4 10-14 f 10 6 10-20 

CO io-13 f 10-7 10-2 0 

1013 -8
 03 f 10 10-21

aMixing ratio. dMeaswed for OH=(v 1). 

bExtrapolated from kq for OH (v 2, 4). eSpencer and Glass (m press). 

cStreit and Johnston (1976). fAnderson (1976). 

Assuming that all other burning processes use approximately the same air/ 

fuel ratio as that of automobile combustion, the total amount of air used 

in all pyrolytic processes is then 3 X 1016 2 yf' globally. The entire atmosphere contains 4.1 X 1021 2 aix (standard temperature and pressure). If 

F-1I and F-1 2 are quantitatively removed in all pyrolytic processes, the 

removal time would be I X i0 5 year. Finally, allowing a factor of 5 for the 

fact that pyrolytic processes usually take place in the source region of 

F-l1 and F-12 (Singh, 1976) the removal time is therefore estimated to 

be approximately 2 X 104 years, which is in good agreement with the 

panel's estimate. 

ION-MOLECULE REACTIONS 

Results of laboratory experiments by Campbell (1976) indicate that a sink 

for F-11 with a lifetime of 70 years (3700 years for F-I 2) may exist in ionmolecule reactions in the troposphere. The 70-year lifetime was obtained 

from measured loss rates of F-11 in air by unidentified reactions initiated 

by y irradiation and is significantly shorter than the lower limit of 03 

years estimated by the NAS panel. The scaling factor for the loss rates of 

F-Il and F-i 2 from Campbell's (1976) experiments to the actual atmospheric condition depends on the interpretation of removal mechanism for 

F-1I and F-12 in the experiments. Campbell (1976) concluded that F-i1 
184 CHLOROFLUOROMETHANE WORKSHOP REPORT 

and F-12 were removed by unspecified terminal ions. However, Fehsenfeld 

and Albritton (1977) have presented an alternative interpretation that 

shows that the loss of F-1I could be accounted fbr by reaction with electrons and leads to a tropospheric lifetime that is roughly 104 longer than 

Campbell's (1976). These two removal mechanisms can be easily tested 

by further experiments as suggested by Fehsenfeld and Albritton (1977). 

ACTIVE REMOVAL INSTRATOSPHERE
 

Photolysis 

Temperature effects on absorption cross sections of CC12F2 and CC13F 

have been determined in the 300- to 210-K temperature range and 200- to 

190-nanometer wavelength region (Chou et al., in press; and Bass and 

Ledford 1976). Stratospheric lifetimes for these molecules are calculated 

by Chou et at. using the temperature-dependent cross sections and 1-D 

transport models. The stratospheric lifetime for CC!2F2 is 9 to 18 percent 

longer, depending on the transport model, than that previously calculated 

using room-temperature cross sections. The corresponding factor for 

CC13F is only 3 to 6 percent. The calculations of Chou et al. also show 

that reduced cross sections result in an increase in steady-state CIX concentration in the middle-upper stratosphere and therefore enhance the 

ozone removal by the ClOx catalyzed chain reactions at steady state. 

Reactions with O(1 D) Atoms 

Rate constants for reactions of 0(' D) atoms with F-11 and F-12 have been 

determined relative to the rate constant of 2.2 ± 0.2 X 10" 0 molecules-' 

cm-a -
 s1 for reaetions of 0('D) atom with N20 (Pitts et al-, 1974). Recently, this reference rate constant has been remeasured to be 1.4 ± 0.1 X - l 0" molecules -' cm+ 3 s -. The rate constants for F-11 and F-12 are 

therefore correspondingly lower by a factor of 1.6, resulting in longer 

removal times for 0('D) reactions. 

CONCLUSION
 

In summary, no significant inactive removal processes have been suggested 

for F-12. For F-11, the combined effects of two possible 70-year sinkssolution in the oceans and ion-molecular reactions-could reduce the 

isteady-state concentration of F-1I by 60 percent. However, because F-i 
ONE-DIMENSIONAL MODELS 185 

contributes only approximately one-third of total-ozone depletion by 

CFM's, a 60-percent decrease in its steady-state concentration will result 

in only approximately 20-percent reduction in predicted ozone depletion. 

EFFECT OF TEMPERATURE FEEDBACK MECHANISM 

In studies of the natural and perturbed stratosphere, most l-D-models do 

not include the effect of temperature coupling (or feedback) mechanisms. 

For local ozone destruction typical of those resulting from the steadystate CFM injection at the 1975 rate, there will be changes in the stratospheric temperature structure. The subsequent effects of these changes on 

the stratospheric chemistry and dynamic phenomena need to be examined. 

Indications are that temperature coupling for the case of CFM injection at 

1975 steady-state rates implies a temperature decrease around 40 kilometers 

of approximately 10 K. Below 30 kilometers, changes are of theorder of 

1 degree. Between 35 and 45 kilometers, these temperature changes result 

in decreased levels of AO3/0 3 due to CFM injection. This is attributable 

to the negative correlation that exists between 03 and T in photochemical 

equilibrium. As a result of the decreased levels of A0 3 /0 3 between 35 

and 45 kilometers,it may be expected that the compensating or "self-heal­

.ng" effect that occurs at the lower stratospheric levels will be reduced. The 

balance between the compensating effect at the upper altitude and the 

reduced "self-healing" effect in the lower stratosphere will determine the 

net effect on (AO3 /0 3 ) due to the inclusion of the temperature coupling. 

Preliminary studies have been conducted in this area; however, resolution 

of this question will follow with more complete results. 

In addition to the impact on the chemical system due to the AT arising 

from -CFM injection, there is also the question of the impact on stratospheric dynamics (see Chapter 5 for additional comment). Under the 

assumptions of a coupled -Dmodel, the stratospheric lapse rate between 

30 and 40 kilometers is reduced by 40 percent due to temperature changes 

associated with CFM injection. Such changes in a parameter important to 

the vertical propagation of planetary waves suggests the possibility of interactions that may significantly impact stratospheric dynamics and, hence, 

transport mechanisms. Results of a study of the sensitivity of stratospheric 

dynamics to radiative transfer mechanisms (Ramanathan and Grose, 1977) 

support this possibility. It is a problem area for future study. 
186 CHLOROFLUOROMETHANE WORKSHOP REPORT 

UNCERTAINTIES IN THE PREDICTIONS OF 1-D STRATOSPHERIC 

MODELS 

INTRODUCTION 

The NAS panel report (1976) discussed the several sources of uncertainty, 

both random and systematic, and found that the two principal causes of 

random error are reaction rates and transport coefficients. They also included in their final predictions of ozone depletion the systematic effects 

of a possible oceanic sink for CFM's. Their estimate of this uncertainty is 

still regarded as current. Uncertainties in transport coefficients are discussed in some detail in Chapter 5. 

REACTION RATES 

Two approaches have been taken to develop reaction-rate uncertainties. 

Chang et al. (NRC Panel-Report, 1976; and NAS Report, 1976) devised a 

sensitivity study in which individual rates are varied in the Lawrence 

Livermore Laboratory model to determine the sensitivity, (r), which is 

defined as follows: 

AlnR r- r=Alnk 

where R is the relative depletion of the ozone column density due to a 

fixed amount of CIX, and k is the reaction rate constant. The uncertainty, 

(u) , due to a particular reaction, (i), is defined as 

u. = r. lnf. 

where f. is the multiplicative uncertainty of the reaction rate. The total 

uncertainty, (U), is 

U 2 = exp Z 1 

The number so obtained has been multiplied by the calculated ozone 

depletion percentage to give what the NAS report (1976) described as 95­

percent confidence limits on the ozone depletion. This method has the 

obvious weakness that the interdependence of different reactions cannot 
ONE-DIMENSIONAL MODELS 187 

be considered. On the other hand, it identifies the specific sources of 

uncertainty and can serve as a guide to kineticists in choosing where to 

concentrate their efforts. 

The NAS panel report (1976) used this method to calculate an overall 

reaction-rate uncertainty of a factor of 2.4 on either side of the central 

value. Their study included uncertainty due to chlorine-nitrate formation 

and five other chlorine reactions, as well as the reaction, OH + HO2. They 

concentrated on the model sensitivity appropriate to small changes about 

the central values and did not quote values averaged over soine range of 

variation. 

This calculation was repeated on the GSFC model at the time of the 

January workshop- Sensitivities to all reactions in the model were considered. Significant sensitivities were found to two more reaction rates 

than the seven originally considered in the NRC report (1976). Thus, the 

overall uncertainties on either side of the central value were found to be 

larger. 

Table 33 shows the sensitivity results obtained with the current evaluation 

of reaction rates (Chapter 1) on the improved GSFC model. This model 

now diffuses odd oxygen, water, molecular hydrogen, and carbon 

monoxide. The overall uncertainties are smaller, and the asymmetry is 

more apparent. 

A second approach to uncertainty analyses has been taken by Stolarski et 

al. (1977),who useda Monte Carlo technique to randomly vary all reaction 

rates independently and simultaneously. The distribution of input reaction rates was assumed to be Gaussian in log-k space with a width adjusted 

to the estimated error in the measured reaction rate. The resulting distributions in calculated quantities were usually Gaussian in appearance, and 

their uncertainty was determined by fitting normal curves to the distributions above and below the central valueFor the current rate set, the Monte Carlo technique gives 95-percent confidence limits of a factor of 1.8 on the high side and 2.86 on the low side 

about a centerline prediction of 16.5-percent ozone-column depletion 

(above 15 kilometers) for an addition of 5.7 ppb Clx corresponding to 

steady-state conditions with the 1975 production rates of F-11 and F-12. 

The high-side uncertainty factor obtained with the sensitivity technique 

agrees quite well with that obtained by the Monte Carlo approach. On the 

low side the sensitivity technique gives a somewhat smaller factor of 2.4, 
188 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Table 33
 

Sensitivities and Uncertainties of CFM-Induced Ozone Depletions
 

Due to Reaction Rates (July Reaction Rate Set) 

GS FC Model 

0to -20 0 to +2u Central 

rl ui r, ui r, U, 

01+03 " CIO+02 0.284 0.079 0235 0U065 0259 0071 

C10+0 --Cl+O 2 0689 0.317 0,594 0.273 0.643 0.296 

CIO+NO - CI+NO2 -0191 .0.132 -0228 -0.157 -0.210 -0.145 

CI + 12 -* HCl + H -0.005 -0.001 -0005 -0002 -0.005 .0.001 

C0+HO - HCI+0 2 -0014 -0.019 -0080 -0110 -0.037 -0.051 

CI+H4 CH HCI+CI13 -0.166 -0.137 -0.194 -0054 .0,194 -0107 

HCI +OH C. CI+H 20 0339 0.078 0.277 0064 0307 0071 

CIO+ N0 2 + M - CION0 2 + M -0252 -0.070 -0356 -0.246 -0300 -0.145 

CIONO 2 +0 -C[O+NO 3 0-006 0004 0.012 0.008 0009 01006 

CIONO + OH - HOCI +NO3 0003 0003 0.009 0.007 0.005 0.005 

0+02 +M - 03+M -0272 -0088 -0287 -0.092 -0-279 -0-090 

0'+0 3 02 +02 -0071 -0033 -004 -0.048 -0.088 -0040 

NO+0 3 - 2N0+02 -0172 -0063 -0.163 -0.060 -0168 -0.062 

OH+03 ' HO +02 -0.335 -0.231 .0.428 -0296 -0.381 -0263 

1102+03 -. OH +0O+0 0058 0.080 0.113 0.156 0086 04119 

NO+ 0 NO + 02 -0.250 -0.035 -0.265 -0037 -0258 -0036 

OiI+0 -"H +02 -04)01 -0.001 -0.013 -0.009 -0.007 .0.005 

H02 +0 - OH+0 2 0.006 0004 0.020 0.014 0.061 0042 

1N0 3 + 0 -: OH+NO3 -0-007 -0031 -0.042 -0.029 -0015 -0040 

N20+O('D) - NO+NO -0.582 -0268 -0694 -0.320 -0631 -0:290 

112-0+O(D) - OH+OH 0.216 0099 0.204 0.094 0213 0098 

H2+O(D) - Oi+H 0.007 0.002 0009 0002 00II 0003 

CH4 + O('D) -+OH +CH3 0027 0.006 0034 0.008 0031 0.007 

0('D)+M -. O(NP)+M 0.382 0123 0372 0.120 0377 0.121 

N+NO - N2 0 0.040 0019 0.046 0021 0.043 0020 

N+O, -. NO+O -0037 -0017 -01135 -0.016 -0037 -0.017 

NO+HO2 - NO2+OH1 0193 0400 0.218 0.151 0.219 0.303 
ONE-DIMENSIONAL MODELS 189 

Table 33 (Continued) 

GSFC Model 

0 to -2u 0 to +2d Central 

iJ' r. ii r, U1 

NO2 +OH+M --- HN0 3 +M 0.433 0.199 0394 0181 0.414 0191 

OH+HNO3 -"NO 3 + H20 -0.073 -0017 -0.095 -0022 -0.083 -0019 

OH+H 20 2 --HO2 +1420 0.050 -0035 -0.065 -0045 0.058 -0040 

OH+ H02 --02 +H20 -0.111 -0.128 -0277 -0319 0.182 -0210 

OH+CH, - CH3 +H20 0015 0007 0020 0009 0017 0.008 

OH +CH20 - HCO +H20 -0039 -0036 -0033 -0030 .0037 .0.034 

HO2 +HO2 "- H202 +02 -0035 -0.049 -0077 -0.107 -0055 -0.076 

CH3 02 +Ho - - CH13001+0 2 0000 0.000 0003 0.012 0001 0004 

nHgh Side* Low Side* Central 

Z-u 2 0625 0876 0.739 

U 1 87 2.40 209 

*The high-side sum iscomposed of positive values of ui from column 0 to +20 and negative values of 

u. from column 0 to -2o. The reverse is true for the low-Side sum. 

which may be attributable to the sensitivity study in adequately accounting 

for the interplay of different reactions as their rates are varied simultaneously. 

Caution should be exercised in the application of these numbers. All of 

them were obtained with the GSFC model, and, although the model has 

compared quite favorably with other models in recent comparisons, the 

sensitivity and Monte Carlo studies have not been conducted on other 

models in their current state. The uncertainty factors should only be applied to the centerline of the model with which they were calculated. 

CONCLUSIONS 

The NAS report (1976) gave a probable range of 2 to 20 percent for the 

steady-state column-ozone depletion due to continued CFM release at 

1973 rates. This range corresponded to their best estimate of the 95-percent 

confidence limits for uncertainties about a centerline value of 7.percent. 

They relied on.one model (that of the Lawrence Livermore Laboratories 
190 CHLOROFLUOROMETHANE WORKSHOP REPORT 

developed by Chang and coworkers) run with as large a variety of input 

parameters as computer resources permitted. The approach taken in this 

report differed in that all of the known active groups in 1-D modeling of 

the stratosphere were asked to participate. Nine groups participated in 

a model comparison in which an attempt was made to run a variety of cases 

of chlorine perturbations on a common set of data. The degree of commonality of the input-data set was greater than that previously achieved 

in attempts to compare models. 

Because of time constraints and differences in model approximations, 

perfect comparison was not achieved, but the overall results indicated 

that, given the same problem and input data, the models of the stratosphere 

will all give essentially the same results. These same modeling groups were 

then asked to give their best estimate of the steady-state ozone depletion 

due to continued release of CFM's at their 1975 rates (which are essentially 

equal to the 1973 rates used by NAS). This is not necessarily meant to be a 

realistic scenario but is meant to be a diagnostic of model results that provides an easy comparison of ozone-destruction efficiency with past estimates such as the NAS report (1976) or with any future-estimates. The 

time-dependent scenarios presented in the NAS report are not expected to 

change significantly, and the comparison of the steady-state values may be 

used to scale any scenarios desired. 

At the time of the January 1977 Workshop at Airlie House, the results of 

of the nine modeling groups for steady-state ozone depletion due to a continuation of the 1975 release rates ranged from 5 to 9 percent when these 

estimates were made with what they believed were the best values of the 

rate constants and other input parameters. This 5- to 9-percent range 

should not be confused with the 2 to 20 percent given by the NAS report 

(1976). The 5- to 9-percent range represents systematic differences in the 

models and systematic differences in the inputs chosen by each group as 

their estimate of the best data. This range does not include all of the uncertainties in rate constants and other input parameters. For instance, if 

a rate has been measured only once, all groups used the same value, even 

though our knowledge of the rate should probably be assigned a fairly 

large uncertainty. Thus, because 5 to 9 percent brackets the centerline 

value given by the NAS report, the conclusion at the time of the January 

Workshop was that, although modeling techniques had improved and input 

data changed, the overall result for column-ozone depletion had come out, 

by a cancellation of changes, to essentially the same number. 

In order to update the model results to account for further changes in rate 

constants, a second, smaller workshop was held at the Goddard Space 
ONE-DIMENSIONAL MODELS 191 

Flight Center in July 1977. The updated rate-constantset is presented and 

discussed in Chapter 1 of this report. The major change was in the rate for 

reaction HO2 + NO -- OH + NO 2 which is now evaluated to be almost 60 

times faster at stratospheric temperatures. This reaction has two direct 

effects on the efficiency of the chlorine catalytic cycle and, hence, on the 

CFM ozone depletion. First, it converts HO2 to OH. The OH is responsible 

for returning inactive HC1 to active Cl. Increasing its rate causes an increase 

in the ratio of active chlorine to inactive chlorine and, hence, in the 

catalytic efficiency. Second, the increase in OH increases the conversion 

ofNO2 (and,thus,NO)to theinactive HNO3 . This decreases the amount of 

NO available to shortcircuit the chlorine catalytic cycle via CIO + NO -> 

Ct + NO2; and it decreases the NO 2 available for formation of inactive 

chlorine nitrate, CLONO 2 . This also leads to an increase in the chlorine 

catalytic efficiency. 

At the July 1977 Workshop, eight of the nine original model groups were 

represented. They all ran their models again for the case of continued CFM 

release at 1975 rates, both with the January 1977 evaluation of the NO + 

HO2 rate constaht and with the new evaluation. This evaluated rate constant 

in Chapter 1 is 8 X 10"2 cm 3 molecule-' sec 1 (i.e., a temperature independent value set equal to Howard and Evenson's (1977) measurement at 

room temperature). Most of the models were run with a slight temperature 

dependence assumed. Of the eight models, six use k = 3 X 10 - 11 c$ 90 / T, 

a value suggested by Howard in his Spring 1977 American Geophysical 

Union Meeting t T presentation. One model (LLL) used 43 X 10-11 e"5° , 

and two models (Ames/R&D and Ames 2-Dimensional) used the evaluated 

number, 8 X 10-

12, with no temperature dependence (see Chapter 5). 

The maximum difference in these rate expressions at 235 K is less than 40 

percent and should be insignificant compared to the factor of 60 change 

from the previous evaluation, but it is a systematic difference between 

models (i.e., assuming that the temperature-independent expression yields 

higher ozone depletions). A few other minor rate differences exist in the 

comparison. Most of the models were run with the January 1977 rate 

evaluation with only NO + HO2 changed. The Lawrence Livermore Laboratory model was run with k 2 =+HO M2 X 10-11 and with rates for HO2 + 

03 closer,butnot equal, to the July evaluations. They also used a CIONO 2 

formation rate with a pressure-dependence similar to the current (July 

1977) evaluation, as did the AmeslR&D models (for the fast NO + HO 2 

only). As is apparent in table 34, these differences appear to be minor 

compared to the effect of the change in the NO + HO2 rate constant. 
192 CHLOROPLUOROMETHANE WORKSHOP REPORT 

In addition to the results shown in table 34, the GSFC model was run with 

the full new rate set exactly as given in Chapter 1,and the resulting depletion was 16.5 percent. An intermediate value for NO + HO2 of 3 X 3 X 

10-11 e-390/T was run on the biechigan model, giving an ozone depletion 

of 8 percent. (Pressure-dependent ClONO2 formation as shown in Chapter 

1was assumed in this run.) As is apparent from table 34, the agreement 

between the 1-D temperature uncoupled models has improved since 

January 1977, even using the same input parameters. The spread is now 

5.8 to 8 percent. This decrease in range should be interpreted as a convergence in the tuning of models to nearly identical inputs and assumptions rather than as an increase in the certainty in the answers. When the 

new NO + HO2 rate is used, all of the models (1-Dtemperature uncoupled) 

showed an increase in the calculated column-ozone depletion, due to continued CFM release at 1975 rates of about a factor of 2 so that the range 

is from 10.8 to 16.5 percent. 

Table 34
 

Ozone Depletion* (in percent)
 

Model Jan. 1977 Rates Fast NO HO2 Factor Difference 

LLL (Lawrence Livermore Laboratories) 7,5 ISO 20
 

I Chang,F Luther,andWDuewer
 

GSFC (NASAiGoddard Space Flight Center) 6.7 130 I 9
 

R Stolarski,D Buflterand$SCbat'dra
 

Dupont 60 122 20
 

P.Jesson, C. Miler,and Fflkm
 

Univenty ofMIchigan 5.8 120 21
 

R Cicerone and S. Liu
 

Ame/RDA (NASA/Ames Research Center/
 

R& D Associates) 3.0 160 20
 

R Whitten antd R Turco
 

AER (Atmospheric and Environmental
 

Researeh, Inc) 59 10.8 18
 

N Sze
 

NOAA/NCAR (NOAA Aeronomy lUbl
 

National Center for Atmospheric Research) 7-0 12,0 1.7
 

P Crutzn andJ.McAeee
 

LaRC Temperature Coupled (NASA/
 

Langley Research Center) 8.7 120 1A
 

L Cahs and J. Nealy
 

Ames 2D (NASA/Ames Research Center) 6.8 163 24
 

R WhittenandW Borucki
 

-Calculated ozone depletion for the January 1977 Workshop rate evaluation and for the update to include 12 3 the fast rate constant for NOO+10 2 ( 8 X 10 cm rnolecule see at 300 K). These ats steady-state 

ozone depletions for continued release of CFM's at their 1975 rates (750,000 metric tons/year). 
ONE-DIMENSIONAL MODELS 193 

The Langley model included a self-consistent calculation of temperature 

by a radiative convective model permitting subsequent feedback -into 

temperature-dependent rate constants. This model shows a somewhat 

larger ozone depletion of 8.7 percent for the January 1977 rate evaluation, 

but is closer to the midvalue for the change to the fast NO + HO2 rate. 

Unfortunately, this modelwas not run in the temperatue-uncoupled mode, 

and no other model was available that had been run for both cases in the 

temperature-coupledmode. Thus,it isdifficult to compare the temperature 

coupled and uncoupled models in any detail to determine if the smaller 

factor difference in the two cases (1.4) is due to the coupling of temperature. Indications from other models'(LLL and GSFO) run for the January 

1977 rates are, that the column-ozone depletion is the same or larger when 

temperature is coupled. The Ames 2-D model results are also shown in 

table 34. They.are essentially the same for latitudinally'averaged columnozone depletion as their 1-) results. 

All of theseresults are the response of 14D models to chlorine perturbations. 

A key question is: How well does the'model response represent the behavior of the atmosphere when it is subjected to such perturbations? One 

way to obtain such an evaluation is to observe the global change in ozone 

taking place after a well-characterized perturbation. The only known 

observation of a catalytic effect of a perturbation on ozone is the comparison ofthe Nimbus-IV BUV datafollowing the August 1972 solar-proton 

event with 2-D model calculations by Heath et al. (submitted for publication, 1976). This was a case of NO, perturbation at high latitude and 

demonstrated that the observations were consistent with catalytic destruction of 0 3 by NOx . 

In general, current evaluations of the relationship between model predictions and the response of the atmosphere rest on the degree of agreement 

between model calculations and concentrations in the current atmosphere. 

For this reason, several sections of this report have been devoted to the 

comparison of theory and field measu'rement. It was generally accepted by 

the working group that comparison of the 14) model results with specific 

single species measurements should not be expected to be better than a 

factor of 2 or 3 because of the assumptions inherent in time and spatial 

averaging and in representing transport in the diffusion approximation. 

Emphasis has been placed on the fact that the best comparisons should be 

those of ratios of constituents whose photochemical exchange times are 

short compared to transport times. For these, the 1I-D model should be 

expected to give results quite close to experiment. Examples of such 

measurements would be O/O 3 , NO/NO2 /0 3 , and C1/C1O. Results from 
194 CHLOROFLUOROMETHANE WORKSHOP REPORT 

such measurements are still sparse. Those that exist are described in 

Chapter 3 and agree reasonably well with models with either value of the 

NO + HO2 rate constant. The ratio of HN3O to NO2 is a somewhat less 

certain test of the model photochemistry in that transport time constants 

are of the same order as chemical time constants. The models should still 

be in general agreement with measurement over an extended altitude range. 

In this case, the one measurement of the ratio is in better agreement with 

1-D model calculations when the slower rate (January 1977 evaluation) for 

NO + HO2 is used. The same is true of the comparison of calculated ozone 

profiles with the ozone profile averaged over many measurements. The 

fast NO + HO2 rate yields larger than observed peak ozone concentration, 

column-ozone content, and tropospheric-ozone concentration. It is not 

clear at this time whether this is due to an inherent inability of the I-D) 

diffusion model to simulate the ozone profile in the transport-dominated 

region to better than about 20 percent, to the fast rate for NO + H0 2 being 

incorrect, or to a needed adjustment in other imperfectly known rate constants. Two-dimensional models also predict ozone columns larger than 

those observed. Thus, although models predict an ozone depletion in the 

10.8- to 16-5-percent range using the current best estimates of rate constants, further changes in the predictions cannot be excluded as both the 

models and the laboratory data are improved. 

A further difficulty encountered in comparing 1-D model results to measurements is the upper stratospheric-ozone profile. Above about 30 kilometers, lime constants for odd-oxygen destruction become sufficiently 

short for one to expect a near chemical equilibrium situation, and the 

model results should be a reasonably accurate representation of the atmosphere in this region. In the 30- to 40-kilometer region, this appears to be 

so when measurements at midlatitudes are compared to model results. 

However, near 50 kilometers altitude, the models appear to give consistently lower values than are measured. A further uncertainty is the 

apparent disagreement between different measurement techniques at this 

altitude. Until this is resolved, it is not clear exactly how large the actual 

ozone concentration is at 50 kilometers and above. It is clear that the 

model results with the current evaluation of rate constants are less than all 

of the measurements. Some attempts have been made to adjust model inputs in order to -lessen this discrepancy. These yield results for ozone 

depletion generally within the ranges already quoted. The adjustments 

made are arbitrary, and it is not clear at this time which, if any, are correct. 

Any of the model results discussed above may be taken as the best answer 

for the centerline value of column-ozone depletion depending on an individual's evaluation of the modeling technique and the input parameters 
ONE-DIMENSIONAL MODELS 195 

used. Even if a best value is chosen there remains the inherent uncertainty in the prediction due to the propagation of the random uncertainties 

in the laboratory or field measurements of inpuit parameters through the 

model calculation. The largest uncertainties seem to be those due to the 

laboratory reaction rate coefficients. A reassessment of their propagation 

by a Monte-Carlo technique (see previous section) led to somewhat smaller 

uncertainties when the current rate evaluation was used. An asymmetry 

in the probability distribution occurs indicating a saturation effect for 

large ozone depletions. The 95-percent confidence levels for propagation 

of random uncertainties (or imprecision) in reaction rates is a factor of 1.8 

on the high side and 2.8 on the low side. Systematic errors due to model 

approximations or missing chemical reactions are not quantitatively 

evaluable. A goal of a measurements program should be to provide data 

which minimizes the probability of such systematic errors. 

In applying these uncertainties to the centerline values, some caveats must 

be made which also apply to the centerline values. These estimates are the 

results of I-D photochemistry models with all of the transport effects 

lumped into one diffusion parameter. Latitude dependent effects will be 

averaged out. More seriously these models do not incorporate feedback 

effects which operate through the radiation and, dynamic fields of the 

atmosphere to modify any chemical perturbations. 
CHAPTER 5
 

MULTIDIMENSIONAL CFM-0 3 MODELING
 

INTRODUCTION
 

The physical structure of the stratosphere is the result of an intricate 

interplay among a large number of processes that, for simplicity, can be 

subdivided into categories of radiation, chemistry, and dynamics (RCD). 

Any physical model of the stratospheric system attempts to incorporate 

at least parts of these components in a self-consistent fashion. Because it 

is not possible to model the stratospheric environment in the laboratory, 

theoretical (mathematical) models must be relied on, subject to observational verification, to simulate stratospheric processes. 

The problems of interactive RCD in their complete generality are complicated beyond current ability to model or to understand. The development 

of a fully generalized model of the atmosphere/ice/ocean RCD system is 

not anticipated in the foreseeable future. Models are already being developed that are significantly more general and self-consistent than any presently available. However, they will still contain substantial simplifications 

of the entire system. 

It is thus clear that current and anticipated models will differ conceptually 

from each other only in their degree of simplification and idealization of 

the real system. However, such models can and do differ drastically in 

their degree of complexity and degree of interactivity among the RCD' 

components. For example, the computational burden required for today's 

most complex models may be more than 107 times that used in a highly 

simplified,but meaningful, one-dimensional (I-D) model. 

Because a generally applicable model of interactive RCD will not be available in the near future, all models must be defined in terms of the problems 

of interest and the types of solutions desired. Inevitably, the model chosen 

is the end product of a series of questions that an investigator must ask 

and answer in order to proceed in an acceptable manner. Examples of 

such questions are: 

" 	What type of approach captures the essence of the processes of 

concern? 

* 	What feedbacks among RCD can be neglected? 

p~g~edn~~ag blan
198 CHLOROFLUOROMETHANE WORKSHOP REPORT 

* 	 Does the model attempt to include processes or feedbacks whose net 

effects are smaller than uncertainties in the major processes? 

* 	What is the weakest link in the conceptual framework? 

* 	 Are the dynamical and radiative simplifications consistent with those 

made in the chemistry? 

* 	 Is there a valid basis for making such decisions? 

* 	 Are the model results to be used for diagnosis or prediction? 

* 	If they are to be used for predictive purposes, how can one distinguish between model sensitivity and an inference that the real 

system will behave similarly? 

* 	 For such predictions, how can a confidence interval on the results 

best be established? 

* 	 As the required confidence interval is reduced, at what point does 

the model become invalid? 

* 	 What are the limitations on computer availability? 

* 	 Are there pressures for achieving "quick" answers? 

The final model chosen is often a very complex end product that reflects 

these objective and subjective considerations. If an investigator chooses to 

work on a rather complex problem emphasizing interactions among the 

various RCD components, he is immediately confronted with the problem 

of how to work with manageable subsets of the more complex processes 

that occur separately in the various RCD components- Almost inevitably, 

the researcher is forced to consider the information learned from noninteractive subset models involving either radiation, chemistry, or dynamics in 

isolation from the other two. Thus, complicated models of the-noninteractive behavior of each of the RCD components are inevitably required for 

developing manageable subsets for use in interactive calculations. 

As an example, complex l-D chemical models are of utmost importance 

for enabling researchers to isolate the important chemical processes from 

the less important and to inspire meaningful approximations and parameterizations. This permits the chemical system to be simplified consistent 
MULTIDIMENSIONAL CFM-C 3 MODELING 199 

with the problem at hand, as well as with the remaining uncertainties in 

the major components. Analogously, important simplifications have been 

made in the problems of dynamics and radiative transfer by employing 

similarly complex, but noninteractive, models. 

However, it should be kept in mind that predictions based on simplified 

noninteractive models (whether one-dimensional or multidimensional) may 

require modification when the full interaction among radiative, chemical, 

and dynamical processes is included. In most instances, these interactive 

processes can be represented properly only in multidimensional models. 

Several examples of important feedback among radiative, chemical, and 

dynamical processes are discussed in the following section. The characteristics of the two-dimensional (2-)) and three-dimensional (3-D)) models 

that have been used in studies of stratospheric ozone are discussed in subsequent sections. The status of these multidimensional atmospheric photochemical models, their current capabilities and limitations, and their possible impact on the chlorofluoromethane 0 (CFM-0 3 ) problem in the 

future are presented in this discussion. Requirements, techniques, and 

problems of the parameterization of transport in 3-D)models are described, 

the observational needs of multidimensional models are presented, and a 

summary is provided. 

RADIATIVE-DYNAMICAL/PHOTOCHEMICAL FEEDBACK 

MECHANISMS 

BACKGROUND
 

The National Academy of Sciences (NAS) report,Halocarbons: Effects on 

Stratospheric Ozone (1976), has identified several sources of uncertainties 

in the estimates for ozone reduction that may result from release of chlorofluoromethanes (CFM's). One such source is feedback mechanisms, which, 

if properly included, could modify the calculated effect of CFM's on 3 

The strongly coupled nature of the dynamical, radiative, and photochemical processes in the stratosphere gives rise to everal direct and indirect 

feedback mechanisms. A few of these feedback mechanisms have been 

included in the state-of-the-art 1-D models, but most have been neglected. 

Although some of the feedbacks can more readily and properly be incorporated in multidimensional models, such models have not progressed very 

far in that direction.. The net effect is that current models are less interactive than they should be for assessment purposes. 
200 CHLOROFLUOROMETHANE WORKSHOP REPORT 

FEEDBACK MECHANISMS 

Stratosphere Radiation * Stratosphere Chemistry 

There are at least two types of feedback in this category. The first con-. 

cerns the effect of a change in stratospheric temperature, T (ST), on 

reaction rates, and the second concerns the effect of the increased transparency of the atmosphere in the ultraviolet (UV) that occurs with a decrease in 0. abundance. The latter coupling gives rise to the "self-healing 

effect" (NAS report, 1976). 

T(ST) . 03 

The absorption of solar radiation by ozone is the main source of heating 

within the stratosphere. As a result, 03 plays a major role in determining 

the stratospheric thermal structure. The stratospheric temperatures, in 

turn, influence the 03 distribution because of the temperature dependence 

of the rates of the various reactions involved in the production and destruction of ozone. This mutual coupling between T(ST) and 03 provides a 

negative feedback in the 1-D models (ie., the inclusion of this feedback 

would reduce the model-estimated value of AO3). Here, AO, refers to 

the percentage reduction in ozone resulting solely from the addition of 

CFM's. The effect of this feedback has been considered by Boughner (1975), Ramanathan and Boughner (1975), and Luther (1976). Ramanathan and Boughner (1975) indicate that this feedback reduces the AO3 calculated otherwise by 4 percent, whereas Luther's model indicates a 10­

percent reduction in AO 3 . Neither of these calculations considers the 

chlorine-nitrate chemistry. Luther's value of 10 percent, perhaps, provides an upper limit on the magnitude of the negative feedback due to the 

T(ST) - 0 3 reduction. 

Caveats 

The l-D models assume the stratosphere to be in strict radiative equilibrium.

Consequently, in the model, at any level in the stratosphere the perturbation in solar heating due.to AO 3 is balanced locally by an equal perturba­ tion in longwave cooling. In the real-world situation, however, a fraction 

of the perturbation in solar heating may be compensated by dynamical 

processes (for example, adiabatic cooling or heating from vertical motion). 

Hence, the negative feedback mechanism of T(ST) # 03, as estimated by 

1-D models, needs to be verified by multidimensional models. 
MULTIDIMENSIONAL CFM-O 3 MODELING 201 

Self-HeallngEffect 

03 destruction at the upper levels in the stratosphere is partially compensated by the increase in 0. at the lower levels due to deeper penetration of 

solar UV radiation. Again, this effect provides a negative feedback. However, this feedback has been considered in almost all of the 1-D models 

and, hence, will not be discussed further here. 

Stratosphere Chemistry f-4T(ST)-4j Atmospheric Transports 

tn the 1-D models, the transport of chemical species by atmospheric motions is incorporated by the use of an eddy diffusion coefficient, K. The 

vertical distribution of K is chosen, to obtain the best agreement with the 

observations for the vertical distribution of the species so that, at best, K 

is appropriate for the present-day stratosphere. However, a global perturbation in 0. would perturb the vertical, as well as the latitudinal, gradient 

in T(STh, which, in turn (as will be explained), can perturb the efficiency 

of the transport processes. This interaction between the gradients in T(ST) 

and the dynamical transports may give rise to several types of feedback 

mechanisms: static stability, latitudinal temperature gradient, and radiative response time. 

Static Stabilit/ Feedback 

The reduction in stratospheric 03 would lead to a reduction in static 

stability. This conclusion can be inferred from the results of Coakley,* 

Geophysical Fluid Dynamics Laboratory staff members (1976), and Ramanathan et al. (1976). The perturbation in static stability could affect the 

mean meridional motions (vertical and latitudinal) and possibly the eddy 

motions. 

Latitudinal Temperature Gradient Feedback 

Even for a AO3 that is uniform with respect to latitude, the variations in 

solar flux and temperature with latitude will affect the latitudinal gradient 

in solar heating and thermal cooling. The perturbation in radiative heating 

would affect the latitudinal temperature gradient, which, in turn, would 

alter the intensity of the zonal winds. Altering the zonal wind intensity 

would affect the propagation of the planetary scale waves within the 

*J. A. Coakley, National Academy of Sciences, private communication, 1976. 
202 CIILOROFLUOROMETHANE WORKSHOP REPORT 

stratosphere, since the propagation of these waves may be very strongly 

dependent on the intensity of zonal winds in the stratosphere (Charney 

and Drazin, 1961; and Schoeberl and Geller, 1976). Since most of the 

contribution to large-scale eddy motions in the stratosphere is from planetary scale waves, it is apparent that a perturbation in 03 may affect the 

large-scale eddy motions within the stratosphere. 

Radiative Response Time Feedback 

The radiative response time, h (the time required for a temperature departure from radiative equilibrium to relax to one-half of its initial value 

or, in an idealized treatment, the inverse of the Newtonian cooling coefficient), is a strong function of temperature, and it is shown in Ramanathan and Grose (1976) that this temperature dependence of h plays 

a substantial role in determining the mean and eddy motions within the 

middle and upper stratosphere. When ozone is reduced, the accompanying 

reduction in T (ST) would reduce t. For example, from figure 8.6 of the 

NAS report (1976), it is seen that a maximum of 40-percent 03 reduction 

occurs at 40 kilometers. (This value includes chlorine-nitrate chemistry.) 

The 1-D models indicate that a 40-percent reduction would cool the 

stratosphere at 40 kilometers by about 10 to 15 K (Coakley;* GFDL staff 

members, 1976; and Ramanathan et al., 1976). From Ramanathan and 

Grose (1976), a 10- to 15-K reduction in T (ST) at 40 kilometers would 

reduce h by approximately 30 to 50 percent. Depending on the vertical 

gradient of the reduction in h, this may either reduce or increase the 

dissipation-of planetary scale waves (i.e., the eddy motions will in general 

be affected). 

To summarize, in view of the mechanisms previously mentioned, it is 

probable that the 03 reduction would affect the transport characteristics 

of the stratosphere. For a 1-D model, the appropriate value of K would 

be modified. As indicated in the NAS report (1976) and in previous 

studies, the computed A03 is sensitive to the value of K. Thus, an initial 

reduction in 03 by the CFM's may perturb the value of K, and this perturbation in Kmayin turn, affect the 03 distribution. 

In the case of multidimensional models, there is generally some attempt to 

model this feedback mechanism. For example, in a 2-D model, the meridional transports may vary with the latitudinal temperature gradient. The 

accuracy with which the feedback mechanism is modeled thus depends on 

*1 A. Coailey, National Academy of Sciences, private communication, 1976. 
MULTIDIMENSIONAL CFM-0 203 3 MODELING 

the reliability of the dynamical parameterizations. In a 3-D model, it is 

realistic to expect increasingty accurate modeling of this feedback 

mechanism (for example, the fundamental equations of motion autonatically incorporate this effect if the radiative heating terms are interactive with 0. perturbations). 

Troposphere Climate Stratosphere Chemistry 

The CFM's have direct and indirect effects on the surface and tropospheric 

temperatures. The direct atmospheric greenhouse effect due to the infrared 

bands of CFM's in the 8- to 12-micrometer region is estimated in the NAS 

report (1976) on the basis of 1-D calculations by Ramanathan (1975) to 

lead to a "steady-state" increase inT of about 0.5 K for a CFM release 

held fixed at 1975 rates. The indirect effect concerns the reduction in 

stratospheric 03 that may result frQm addition of CFM's and the subsequent effect of 03 reduction on the surface and tropospheric temperatures. This indirect effect, through 03 changes and their subsequent 

effect on the greenhouse mechanism, is probably smaller than the direct 

CFM greenhouse effect. These changes in the tropospheric temperatures 

may subsequently affect the stratospheric chemistry by modifying transport. 

Tropopause T-Stratosphere H20 Feedback 

The possibility of this feedback for the ozone reduction problem has been 

mentioned by Liu et al. (1976). As discussed in the NAS report (1976), 

this feedback is still in the hypothesis stage and has not been verified by 

model studies. The concern for this feedback arises from the Mastenbrook 

(1971) study that indicates that the mixing ratio of H20 in the stratosphere 

is approximately equal to the saturation mixing ratio of H.0 at the tropical tropopause temperature. This had led to the suggestion that the H2 0 in 

the stratosphere is determined by the tropopause temperature of the tropics. According to the NAS report (1976), the direct and indirect effects 

of the CFM's may cause an increase of about 2 to 3 K in the tropopause 

temperature, which would increase the saturation mixing ratio of H20 by 

a factor of 1.4 to 1.7. Liu et al. (1976) conclude that increasing stratospheric H.O significantly enhances the 03 reduction caused by a given 

amount of CFM's. For example, Liu et al.'s (1976) result indicates that 

doubling the present-day H0 concentration in the stratosphere nearly 

doubles the ozone reduction caused by the addition of CFM's. Thus, this 

coupling provides a positive feedback. 
204 CHLOROFLUOROMETfNE WORKSHOP REPORT 

Surface Temperature-Cloud Coupling 

As mentioned previously, the direct climatic effects of CFM's is an increase 

in the global surface temperature,T s . It is conceivable that the increase 

inTsmay alter the global cloud cover (Schneider and Dickinson, 1974). A 

perturbation in the cloud cover would perturb the solar radiation backscattered by the clouds and, consequently, would affect the photolysis of the 

trace species. To estimate the magnitude of this feedback, it is necessary 

first to know the nature of the T!-cloud cover feedback, but, unfortunately, even the sign of this feedback is not known. Cess (1976) estimates 

from observed cloud climatology that the maximum cloud cover change 

may be approximately 0.02, from the unperturbed value of 0.4 to 0.5 fractional cloud cover, for a change inT of 1 K. Although there are presently 

no model calculations to indicate the effect on 03 of such a change in 

cloud cover, available 1-D models could be used for appropriate calculations. It seems unlikely that this effect is large. 

Tropospheric Latitudinal Temperature Gradient-Stratospheric Transport 

Coupling 

The 1-D greenhouse calculations indicate only the increase in the globally 

averaged surface temperature. However, it is known that the response of 

the surface temperature to perturbations may differ significantly from one 

latitude to another- For example, the recent 3-D climate model results of 

Manabe and Wetherald (1975) for the CO2 -climate problem indicate that 

the simulated increases in surface temperature and lower tropospheric 

temperatures are significantly larger (by a factor of 3 to 4) in the polar 

latitudes than in the equatorial latitudes. A perturbation in the latitudinal 

temperature gradient would perturb the tropospheric motions, which, in 

turn, may perturb the stratospheric motions, because mechanical forcing 

by tropospheric motions is one of the primary energy sources for stratospheric motions. The CFM-climate problem is analagous to the CO2 -

climate problem, and, hence, it appears to be possible that the effect of 

CFM's on thetropospheric climate can perturb the transport characteristics 

of the stratosphere. 

Caveats 

Finally, the results mentioned here for the increase ii the surface and 

tropopause temperatures are based on I -D radiative-convective model calculations. The limitations on the application of such calculations to the 
MULTIDIMENSIONAL CFM-O3 MODELING 205 

real world have been discussed in detail in Chapter 6 of the NAS report 

(1976) and are therefore not repeated here. 

CONCLUDING REMARKS 

The preceding discussions illustrate the complex coupling and feedback 

mechanisms among various physical processes and between the troposphere 

and stratosphere. The mechanisms discussed previously are only some of 

many that may conceivably occur. It is necessary to have quantitative estimates of the different potential feedback effects to determine which are 

important and which are negligible. For those discussed previously, it is 

possible to be sermiquantitative in only a few instances. 

Quantitative evaluation of feedback mechanisms will generally depend on 

the development of appropriate fully interactive models. In most cases, 

this will require use of 3-D models, although, once-a feedback mechanism 

is adequately understood, it might be possible to include it in models of 

lower dimension by means of appropriate parameterizations. The task of 

making models fully interactive includes the consideration of all parameterized processes as well as processes that are explicitly modeled. Threedimensional models are particularly appropriate for investigating feedbacks 

because they explicitly include many interactions; however, it should be 

emphasized that many processes are parameterized in present 3-D models. 

In general, it is essential that careful analysis of the process be made in 

addition to interactive simulation experiments. 

TWO-DIMENSIONAL MODELING 

The 2-D model is usually formulated as a phenomenological model similar 

in general concept to the l-D model. However, it includes the capability 

of describing atmospheric latitudinal gradients and seasonal behavior by 

accounting for the important time-dependent meridional transport in the 

atmosphere, together with the seasonal and latitudinal variations of the 

solar zenith angle. Since the models are phenomenological and are not 

developed completely from first principles, it is essential to develop and 

substantiate the models by comparison with observations. It should be 

remembered that, although development from first principles is always 

desirable, it is only a prioil necessarily better than a phenomenological 

model if the development includes a realistic description and calculation of 

all pertinent physics needed to solve the governing equations. As a result, 

in either, case, extensive testing of the model against actual observations is 
206 CHLOROFLUOROMETHANE WORKSHOP REPORT 

needed to verify the results. These tests should encompass at least four 

general areas: chemical system (reactions and associated rates), atmospheric transport, combined chemistry and transport, and boundary conditions. It should be emphasized that information necessary for a comprehensive testing of a 2-D model is at present limited; therefore, verification 

of certain aspects and/or regions of the model are also limited. Some of 

these aspects are pointed out in the following discussion. 

Because the purpose of this subsection is to briefly review the current 

status of 2-1) modeling, the history of these models isnot included, and 

only the latest available information on each model has been used. A brief 

description of some of the pertinent aspects of the models is given, and the 

general capabilities1 limitations, and use of 2-D models in atmospheric 

problems are then discussed. For a complete description of each model, 

the reader is referred to the indicated references. 

Two-dimensional atmospheric photochenical models have been developed 

by Brasseur and Bertin (1976), Borucki et al. (1976), COMESA (1975), 

Crutzen (1975), Harwood and Pyle (1975), Prin et al. (1975), Turner et 

al. (1976), Vupputuri (1976a), Whitten et al. (1977), and Widhopf 

(1975 and 1976). Because Prin et al. (1975) developed their model as an 

intermediate model to be used as a source of NO, distributions for their 

3-1) model and used measured distributions of ozone rather than calculating the ozone concentrations, there is no further discussion of this model 

in this section. 

TYPES OF MODELS 

For purposes of discussion, the current 2-D models can be divided into 

two classifications. In the strictly photochemical model,the hydrodynaniic 

characteristics of the atmosphere (man density, temperature, meridional 

wind field, and eddy-diffusion coefficients) are prescribed and the species 

conservation equations solved to determine the time-dependent distribution of the atmospheric trace constituents of interest. Atmospheric 

transport by mean meridional winds and eddies are included in the models. 

The eddy mixing is developed by using a mixing-length hypothesis developed by Reed and German (1965). Because the transport parameters are 

obtained directly or indirectly from observations, they are not necessarily 

unique and must be fully tested to determine their usefulness. Due to 

uncoupling of the chemistry and transport, these models are designed (as 

are I-D models) to investigate small changes in the chemical structure of 

the atmosphere, which are, in turn, implicitly assumed to have a negli­

-gible effect on the transport and temperature distributions. 
MULTIDIMENSIONAL CFM-0 3 MODELING 207 

In the second classification, an atmospheric radiation model is included, 

and the temperature and mean circulation are calculated as part of the 

solution. Eddy diffusivities must be prescribed in either approach. The 

models of Harwood and Pyle (1975), Turner et al. (1976), and Vupputuri 

(1976b) fall in this second category. Except for Vupputuri's (1976) model, 

the radiation-coupled models are in too primitive a stage to assess their 

full capabilities to predict the species distribution in the natural atmosphere throughout the year. This second type of model presents the hope 

of a more coupled and consistent evaluation of the changes in atmospheric 

properties that occur as a result of the introduction of pollutants, since 

the mean circulation and temperature are coupled with the calculation of 

the species concentration. It should be remembered that the coupling 

with the transport is not complete since these models are still dependent 

on a prescription of an important governing parameter (namely, the eddydiffusion coefficients). A realistic evaluation of the model capabilities 

must await the completion of more work by the respective workers. 

All of the models simulate seasonal changes in the meridional plane 

(essentially pole to pole, surface to stratopause) except for Vupputuri's 

(1976b), who considered a steady-state seasonal model (winter, summer) 

from 10 to 50 kilometers. Thus, interpretation of his results should be 

made with this inherent limitation in mind. 

CHEMICAL MODEL 

With regard to the previously outlined tests, all of the 2-D models have, 

with some variation, essentially relied on the results of 1I) evaluations to 

ascertain which chemical model to use. Thus, the chemical systems are 

basically the same, considering the typical O., NOX, and HOx systems 

outlined in the Climatic Impa~t Assessment Program (CAP, 1975), with 

the following exceptions. Vupputuri and Widhopf do not consider NO3 

and N2 0 5 . Vupputuri (1976b), Brasseur and Bertin, and Whitten et at. 

have considered ClOX. COMESA, Crutzen (1975), and Widhopf (1976), include the smog chemistry cycle arising from the oxidation of methane. Presently, only Crutzen (1975) includes the calculation of water vapor; other 

modelers prescribe this distribution. Basically, the reaction rates used are 

those recommended by the National Bureau of Standards review; however, 

there are some important differences. One of the more important of these 

is the rate at which the reaction HO2 + OH -+ H2 + 02 proceeds, and 

others are outlined in the NAS report (1975). The results of each individual model have to be evaluated with these differences in reaction rates 

and their subsequent effects in mind. 
208 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Current differences between measurements ofvarious trace species and 1-1) 

calculations can be more fully evaluated byperforming tests using 2-D 

models since horizontal transport and seasonal variations would be included 

in the evaluation. These tests are necessary to more fully evaluate the 

chemical systems used in the atmospheric models and to affirm the choices 

made using l-D models. These evaluations depend on the availability of 

more measurements of the important atmospheric'species for adequately 

determining their seasonal and latitudinal variations. Also, since the models 

represent average concentrations in either space (zonal) or time (diurnal), 

numerous measurements are needed in order to make the comparisons of 

the calculated distributions with data more meaningful. Above all, it 

should be remembered that the model is as good as the chemistry and 

transport utilized. 

ATMOSPHERIC TRANSPORT PRESCRIPTIONS FOR 2-D MODELS 

The eddy-diffusion coefficients used in each model are different, using the 

turbulent diffusion coefficients of Luther (1973), COMESA, Vupputuri 

(1976b), improvements of Luther's coefficients (Widhopf, 1976), ones 

derived by imposing a match between model ozone calculations and observations (Crutzen, 1975; Brasseur and Bertin, 1976; and Whitten etal., 

1977) or ones based on observed distributions of 0., NO1 , and 14C 

(Borucki et al.). The mean wind fields used by various modelers include 

those derived by Louis et al. (1974), Crutzen, (1975), and Widhopf (1976), 

that obtained from 3-D model calculations (Brasseur and Bertin, 1976), a 

rather arbitrary prescription (COMESA), mean winds calculated in the 

model (Vupputuri, 1976b), or mean winds based partially on observations 

and partially on 3-D model calculations (Borucki et al, 1976). 

The simulation of the variation of water vapor in the atmosphere, as well 

as the dispersion of volcanic debris, has also been used by Crutzen (1975) 

to develop his transport. Water vapor is very sensitive to the atmospheric 

exchange between the troposphere and stratosphere and is useful in testing 

model transport; however, data on water-vapor concentrations in the 

stratosphere are limited. 

Regarding the testing of the atmospheric transport used in models, 

COMESA and'Widhopf have tested their models against the dispersion of 

radioactive nublear debris. COMESA finds the downward transport too 

fast and the poleward transport too small using Luther (1973) eddy diffusivities. These results are in general agreement with those obtained by 
MULTIDIMENSIONAL CFM-0 209 3 MODELING 

Widhopf (1975) using the Luther (1973) coefficients and Louis circulation. Glatt and Widhopf (1976) improved the coefficients of Luther 

(1973) using as a guide Luther's (1973) unsmoothed (raw) data and the 

carbon-14 measurements. This approach yielded reasonable agreement 

of the model calculations with the time-dependent dispersion of carbon-14 

throughout the northern latitudes. In addition, good agreement was observed with the dispersion of tungsten-185 in the equatorial regions, and 

relatively good agreement was obtained with the dispersion of Zr-95 in 

which particle settling was included to account for this important effect 

above 20 kilometers. 

Fabian and Libby (1973) tested some combinations of mean circulation 

and eddy-diffusion coefficients and found the best relative agreement with 

Zr-95 data using the turbulent diffusion coefficients of Luther (1973) 

and the circulation of Reed and German (1965). In addition, they obtained some estimates of atmospheric residence times. 

Tests of model simulation of the dispersion of inert tracers are essentially 

independent of a knowledge of the photochemistry of the atmosphere and 

and thus are desirable when applicable. This is especially important for 

pollution problems, in which not only a knowledge of the chemistry is 

important, but also a knowledge of the local residence times. In this 

regard, tests using ozone, N20, CH4 , or any other chemical species are 

dependent on a knowledge of the important chemical sources or sinks for 

that species and may not be an entirely independent test of the model 

transport. It should be pointed oit, however, that tests using radioactive 

debris data are subject to the following uncertainties: accuracy of the data, 

incompleteness of the measurements'in space and time, modeling of the 

rainout/washout sink in the troposphere, and particle settling. These 

questions make comparisons with inert tracer data somewhat subjective, 

and care must be taken in the interpretation of the results. 

It should also be noted that, due to measurement platform limitations, 

nuclear debris data are limited to below 25 kilometers and essentially the 

Northern Hemisphere. This makes model verification in the Southern 

Hemisphere and above 25 kilometers difficult. The amount of data on the 

distributions of chemical species used to develop the model transport is 

also scarce in the Southern Hemisphere and at high altitudes. As a result, 

these are definite areas in which more information is needed to evaluate/ 

develop atmospheric models. 
210 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Only by testing against various specie distributions in the atmosphere 

(ozone, N2 0, HN0 3 , NO., CO, and H20) and various inert tracers can a 

model be adequately verified. This can be done adequately as more measurements of all these constituents become available. 

CHEMISTRY AND TRANSPORT 

The referenced 2-D models have been used to simulate the distribution of 

trace species in the natural atmosphere. In one sense, this can be looked 

upon as another test of the combined model chemistry and transport when 

compared with data and one which provides information on species concentrations in regions in which measurements have not been made. Each 

of the models generally shows relatively good agreement with the observations of most atmospheric species. However, these measurements are sparse 

both in latitude and in time during the year. As with I-D models, the 

difficulty in any comparison is that the measurements are made at one 

place at one specific time of the day, whereas the calculations are diurnal 

averages and represent "statistically averaged" concentrations. Most of 

the models are able to reproduce the seasonal and latitudinal variation of 

the ozone column, some more successfully than others. Those modelers 

who used ozone as a guide to develop their transport naturally show good 

agreement with ozone measurements. Detailed information on the calculated monthly variation of the ozone column has not been presented by 

most of the modelers; however, the distributions shown for particular 

seasons/months are in reasonable agreement with measurements. Brasseur 

and Bertin (1976), Crutzen (1975), and Widhopf (1976) have presented these 

distributions, and their models satisfactorily simulate the monthly variation 

of the ozone column in the Northern Hemisphere and show adequate 

agreement in the Southern Hemisphere. COMESA has also presented such 

comparisons; however, their results do not show good agreement with data. 

BOUNDARY CONDITIONS 

Boundary conditions used in the models are varied, and, in general, few 

sensitivity studies have been performed to study the effect of changing 

them in any given model. More information regarding the flux of various 

constituents and/or their concentration is needed in order to improve the 

boundary conditions in any photochemical model. Information is also 

needed with regard to the latitudinal and seasonal variations at the boundaries. This aspect of the modeling, as well as sensitivity studies of the 

chemical system (reactions and rates) used in the models, should be given 

more attention in the future. 
MULTIDIMENSIONAL CFM.0 211 3 MODELING 

APPLI CATIONS 

Applications of the 2-D models have been used in the analysis of the effect 

on ozone of NO emitted by aircraft and produced by solar-proton events. 

Various modelers (Crutzen, 1975; Vupputuri, 1976b; and Widhopf et al., 

1976) have investigated the effect of NO, emissions of estimated future 

fleets of supersonic aircraft during the CIAP. Their globally averaged resuits were in general agreement with the spectrum of 1-D results; however, 

they showed a large latitudinal effect as a result of the projected flight 

corridors. A large seasonal variation was also obtained by Crutzen (1975). 

More recently, Borucki et al. (1976) have performed similar calculations 

with their model, obtaining results in general agreement with other 2-D 

and 3-1) model results. 

More recently, Hildalgo and Crutzen (1976) and Widhopf et al. (1977) 

have applied their models to the calculation of the effect of realistic 

estimates of future supersonic and subsonic fleets. In both studies, it was 

found that ozone was produced in the troposphere by the subsonic aircraft NOx emissions, whereas the supersonic aircraft emissions catalytically 

destroyed ozone. The net difference was a slight (less than 0.6 percent) 

northern hemispherically averaged increase in ozone with significant 

latitudinal and seasonal variations. As stated previously, these models include the smog chemistry cycle initiated by methane oxidation by OH. 

Crutzen (1975) has also applied his model to the calculation of the effect 

of the 1972 solar proton event, calculating an ozone reduction in agreement with satellite observations. This is very important since it is the 

only direct verification of the theory of catalytic ozone destruction by 

NOx It is a good example of a pertinent contribution of a 2-I) model because this effect took place in the northern polar region and, hence, could 

not be adequately treated by a 1-D model. 

To date, three 2-I) models have been applied to the CFM problem.(Vupputuri, 1976; Brasseur and Bertin, 1976; and Whitten et al., 1977). Vupputuri (1976) applied his 2-D steady-state model, including CIO, and 

CIONO 2 using constant 1973 injection rates of CF2 Cl and CFC1l and ran 

the solution to a steady state. Unfortunately, this type of model does not 

have all of the advantages of a time-dependent model, but it.does provide 

some information. Vupputuri (1976) used the reaction rates tabulated in 

the NAS report (1976) and found a globally averaged decrease Of ozone of 

approximately 8 percent- The maximum reduction was about 10 percent 
212 CHLOROFLUOROMETHANE WORKSHOP REPORT 

in the Northern Hemisphere (winter) and decreased to 6 percent in the 

Southern Hemisphere (summer). The net effect of including ClONO2 was 

a factor of 1.4 decrease in the reduction of ozone for the same injection 

rate. 

Brasseur and Bertin (1977) also applied a steady-state version of their 

model to the CFM problem. These calculations are very preliminary and 

are to be superseded by a time-dependent calculation in the next few 

months. 

Whitten et al. (1977) have run their time-dependent 2-D model using the 

CFM release rates of Howard and Hanchett (1975) extrapolated to 1976 

and held constant thereafter; the reaction rates are those recommended in 

this volume, except that the slow rate for reaction 31 (1.5 X 1011 exp - (-1 00/T) cm3 see) ) was also used for a comparison run. Chlorine nitrate 

was included in the model. 

The use of the recommended rate coeffidient for reaction 31 leads to a 

globally averaged rise in the predicted ozone-column density of about 15 

percent. Widhopf obtained a similar result with his 2-D model. As in the 

case of 1-D model calculations, predicted OH distributions come into better 

agreement with Anderson's (1976) measurements when the recommended 

coefficient for reaction 31 is used. 

The predicted 1990 asymptotic chlorine mixing fraction was 1.0 ppbv. 

The corresponding globally and seasonally averaged ozone reductions were 

about 1.2 percent (using the slow rate coefficient for reaction 31 and 

ignoring formation reaction 50 and about 2.9 percent (using the recommended rate coefficient for reactions 31 and 50). Investigation of the 

relative ozone reduction as a function of altitude showed that it was 

qualitatively similar above 35 kilometers for both values of rate coefficient 31. However, at lower altitudes, the behavior of the relative ozone 

reduction was very different. For the low value of rate coefficient 31, 

an increase in ozone was predicted because of the role of chlorine nitrate. 

However, when the recommendedrate coefficient for reaction 31 was used, 

the predicted ozone below 30 kilometers was found to decreaseas aresult 

of the nearly fourfold increase in chlorine oxide; only at low latitudes did 

the predicted ozone experience a slight increase. At midlatitudes, these 

results are in very good agreement with those obtained with 1-D models. 

As discussed'previously, the use of a time-dependent 2-D model to investigate the CFM pollution problem would provide information regarding 
MULTIDIMENSIONAL CFM-0 3 MODELING 213 

seasonal and latitudinal variations, interhemispheric mixing, and the importance of variable upward and downward transport. These may be important factors to consider in any comprehensive evaluation of this 

problem. However, any such calculation using a time-dependent model 

would require about 10 to 20 hours of CDC-7600 time to simulate the long 

periods of time involved in calculating some of the scenarios considered 

by the NAS report (1976). This may be necessary, however, in order to 

obtain a multidimensional picture of the 1-D results. Various modelers are 

currently working on ways to reduce the amount of computer time 

required. 

Another point of interest would be the investigation of the indications that 

there is a 20-percent variation in CFM concentrations between the Northern and Southern Hemispheres and the test of whether multidimensional 

models can explain/reproduce these measurements. 

SUMMARY 

Work by various modelers has shown that a 2-D model can be developed 

that includes a more realistic picture of the atmospheric transport than 

that of a 1-D model (namely, horizontal transport, latitudinal gradients, 

and seasonal changes). This can be done with a reasonable expenditure of 

computer time using a comprehensive chemical model. These phenomena 

are important because latitudinal gradients are large, and, thus, the upward 

and downward transport of a constituent need not necessarily be the same 

or latitudinally uniform as in a 1-) model. This could be important in 

some pollution studies in which a constituent moves upward at a particular latitude/season and may move downward at another latitude/season 

at a different rate. 

Two-dimensional atmospheric models have been able to adequately simulate the latitudinal/seasonal distributions of many atmospheric trace 

species, including ozone, NOy, and HNO3 . They have revealed significant 

latitudinal and seasonal variations in the distributions of various other 

atmospheric trace constituents. Inherent in these latitudinal and seasonal 

variations is the important effect of the corresponding variation of the 

local solar flux, which can be accounted for in a 2-D model. When applied 

to NO. pollution problems, their globally averaged results are in agreement 

with the range of corresponding 1-D results;however, they also show large 

latitudinal and seasonal variations about this average. They also provide 

the ability to study interhemispheric mixing and hemispheric differences 

and the ability to study atmospheric problems that are basically multidimensional. All of these aspects should be included in any comprehensive 
214 CHLOROFLUOROMETHANE WORKSHOP REPORT 

evaluation of the effect of the introduction of a pollutant into the atmosphere. A 2-D model can be used to obtain estimates of these effects, using 

at least one order of magnitude less computer time than that needed for 

a 3-D computation. 

However, 2-D models are still in a development state in which more data 

are needed to verify them. They must also be developed further to include 

scattering effects, more coupling between the dynamics and chemistry, 

more accurate rainout/washout modeling, better boundary conditions, and 

more sensitivity studies. The results of these models will enhance current 

understanding of atmospheric problems. 

THREE-DIMENSIONAL MODELING 

A 3-D model has the potential to provide the only physically realistic 

treatment of the transport of trace species in the atmosphere. Calculations 

(Mahlman, 1975) indicate that there is no correlation between the rate of 

transport of a trace species and the instantaneous gradient of the species. 

These model results are consistent with the implications of observational 

numerical and theoretical results which show that, under normal conditions, there is a near balance between the horizontal and vertical transport 

by eddies and the transport by the mean meridional circulation. In fact, 

even during transient events such as sudden warmings, the net transport is 

only a small residual difference between opposinglarge contributions from 

eddies and the mean meridional circulation. However, the 3-D tracer model 

of Mahiman (1973) shows that this strong compensation can break down 

significantly at times, particularly following seasonal circulation transitions. 

Theory indicates that net transport depends primarily on distributions of 

thermal and viscous damping and on wave transience, rather than on 

gradients of mean quantities. However, in l-D and 2-D models, transport 

must be parameterized as proportional to some function of the mean concentrations. In 1-D and 2-D models, such a correlation is assumed to exist 

in the mean. 

One-dimensional and two-dimensional models also possess the limitation 

that transport parameters are assumed to be independent of the distribution of sources and sinks of trace species in the atmosphere and are not 

*permitted to change in response to anticipated dynamical changes. Therefore, these models are limited to treating small local changes. It should be 

noted in figure 8.6 of the NAS report (1976) that the ozone depletion 

ratio associated with chlorocarbons ispredicted to maximize at an altitude 
MULTIDIMENSIONAL CFM-0 3 MODELING 215 

of approximately 40 kilometers. At this altitude, stratospheric motions are 

driven at least partially, by the absorption of solar ultraviolet radiation 

by ozone. Significant changes of the transport rate of trace species due to 

ozone changes at 40 kilometers should therefore be anticipated even if the 

total columnar ozone loss should not exceed 10 percent (as previously 

discussed in this chapter). The temperature structure of the stratosphere 

could also change, with resulting changes in the rates of the various 

chemical reactions. A potential change of the transport rate of trace 

species illustrates the possible limitation of most 1-D and 2-D models of 

the atmosphere (which are incapable of responding to such a variatibn of 

stratospheric climate) and an inherent advantage of most 3-D models. 

The ability to simulate atmospheric transport is not the only way in which 

3-D models should be capable of providing a more realistic simulation of 

the atmosphere and of estimating the effects of chlorocarbons. A 3-D 

model can treat atmospheric inhomogeneity in both space and time. 

Stratospheric chemistry contains many reactions that introduce into the 

numerical models terms of the form, kAB [A] [B], where [A] and [B] 

are species concentrations, and kABwis areaction rate. This nonlinear term 

is usually spatially and temporally averaged in the models, and deviations 

of each species and rate constant from its mean value and correlations are 

often ignored. The most obvious deficiency in this approach is associated 

with diurnal variations, because many of the species of stratospheric interest are absent at night. One-dimensional models that incorporate diurnal 

variations now exist (e.g., Turco and Whitten, 1977), and techniques for 

incorporating this effect into multidimensional models are being implemented. 

On the other hand, spatial correlations should be simulated using multidimensional models. Reactions that have an activation energy exceeding 

1000 kilocalories could lead to significant correlations, particularly in view 

of the substantial (approximately 50 K) temperature variations in stratospheric warmings. Observations of NO and C1O (see chapter 3) suggest 

variability of more than a factor of 2, and theoretical considerations and 

NO2 observations by Noxon (1976) suggest significant latitudinal gradients of some species that are probably associated with conversion to other 

species such as N2 O. and C1NO at high latitudes in winter. Assessments 

of the effects of latitudinal gradients on these correlations can be made 

with a 2-D model, but a full assessment of these effects to include local inhomogeneity can be simulated only in a 3-D model. 
216 CHLOROFLUOROMETHANE WORKSHOP REPORT 

Furthermore, as pointed out by Martin and Stewart (1976), the variability 

of atmospheric molecular oxygen concentrations and the strength of the 

Shumann-Runge bands lead to a moan production rate of odd oxygen 

which is different from that obtained by calculating the photodissociation 

rate using a mean profile for molecular oxygen. Similar errors may be 

introduced in other photodissociation rates that depend on absorption in 

the Shumann-Runge bands (e.g., H20) and by using mean temperatures 

for strongly temperature-dependent reaction rates (e.g., the destruction of 

odd oxygen by the 0 + 03 reaction). 

Because 3-D models attempt to simulate atmospheric variability, they 

should provide a more realistic test of the validity of the chlorocarbon/ 

ozone theory from the limited number of observations of chlorine species 

that will be made during the next 2 years. 

Three-dimensional models of the stratosphere have been under development since the time of Hunt and Manabe (1968)- Most of these models 

have been concentrated on representing atmospheric motions, and only 

recently has much consideration been given to the techniques for including 

a detailed chemistry. These models consume large amounts of computer 

time on the fastest available machines (eg., several seconds to 1hour per 

model day), and, thus, the fundamental problem of constructing such a 

model is in estimating the spatial and temporal scales-that must be retained 

explicitly in order to simulate the stratosphere. 

A 3-D model that is sufficiently developed to have the potential for including the minimum required chemistry for the chlorocarbon problem during 

the next 2 years isthat of Cunnold et al. (1975). Large modelssuch as that 

of Malilman (private communication, 1977) will probably require 5years or 

more to ilcorporate fully consistent chemistry. The current version of the 

model of Cunnold et al. (1975) is being revised to provide a more authentic 

simulation of the atmosphere. At the present time, these models only 

partially incorporate radiative feedbacks. For example, the model of 

Cunnold et al. (1975) contains a full treatment of the absorption of solar 

ultraviolet radiation by ozone, but it parameterizes the absorption of 

infrared radiation using a Newtonian cooling approximation with a 

pressure-dependent coefficient. As previously mentioned, substantial 

changes in this coefficient may occur in response to the accumulation of 

fluorocarbons. Thus, more realistic treatments of infrared radiation should 

be incorporated into these models. 

It appears that existing 3-D models that include six planetary waves do not 

provide sufficient horizontal resolution for adequately simulating the 
MULTIDIMENSIONAL CFM-O3 MODELING 217 

the troposphere/stratosphere exchange process (see Mahlman, 1975). 

Only one-fourth to one-third of the observed downward transport of ozone 

of 25 tons per second has been obtained in a model with this resolution. 

Atmospheric budgets of many chemical species, including chlorocarbons, 

are controlled bythe transport through the tropopause, which, in the mean, 

occurs in 1 to 2 years. It is therefore important that this vertical transport 

process be simulated adequately. The amount of parameterized transport 

required in a 3-D model depends critically on its spatial resolution. (See 

the following section.) As an initial effort toward achieving most of the 

transport explicitly, it is planned to increase the horizontal resolution of 

the Cunnold et al. (1975) model to at least twelve planetary waves. In 

conjunction with these calculations, it is desirable that more atmospheric 

observations be made, placing emphasis on the nature of upward transport 

from the troposphere into the middle stratosphere. 

Although it has been shown that the spatial resolution necessary for 

simulating stratospheric motions is less than that for the troposphere 

(Washington, 1972), there remains an unresolved question concerning the 

resolution that is needed in the troposphere in order to simulate adequately 

the large-scale motions therein that are responsible for the temperature 

structure of the lower stratosphere. It is assumed that a resolution of 

twelve planetary waves will also be sufficient for this purpose. 

Thus, increased spatial resolution should improve the authenticity of those 

3-D models that incorporate a few chemical reactions. In addition, as 

recently shown by Mahlman (private communication, 1977), increased 

vertical resolution may be necessary to overcome the common discrepancy 

that is seen in model results at high latitudes in winter, in which the temperatures at the upper levels of the models are too low by 20 to S0 K. As a 

result, simulated polar night jets continue to increase in height up to the 

upper boundary. This feature of the models may be responsible for the 

achievement of only limited success in the simulation of sudden stratospheric warmings. For example, although substantial temperature oscillations in the stratosphere of a magnitude similar to observations have been 

obtained within a 2-week period, the calculated zonal structure of the 

stratosphere has prevented these oscillations from reversing the zonal wind 

and Equator-to-pole temperature gradient. Proper simulation of these 

features is important because substantial transports of both heat and trace 

species occur during both major and minor stratospheric warmings (e.g., 

Newell et al., 1974; Hartmann, 1977; and Madden, 1975). 
218 CHLOROFLUOROMETHANE WORKSHOP REPORT 

It should be possible to incorporate a set of chemical reactions that is 

consistent with the model resolution into a 3-) model during the next 

several years. The addition of chemical species to such a model must 

include a careful assessment of chemical time constants. Species having 

time constants that are short compared with the model time step (about 

I hour) should be calculated under the assumption of photochemical 

equilibrium. Each longer time-constant species that is permitted to be 

transported by the predicted motion field adds approximately 30-percent 

additional computation time and must therefore be carefully selected. 

This type of model can be integrated for approximately 3 years, after 

which dynamical equilibrium of a species such as ozone is attained. However, the ultimate buildup of chlorocarbonsin the atmosphere is determined 

by a balance between input and loss and may take 50 to 100 years to be 

attained (NAS report, 1976). Thus, calculations that attempt to simulate 

the continuous accumulation of chlorocarbons in the atmosphere are not 

possible with a 3-D model; moreover, special techniques that short-circuit 

the accumulation process must be employed to determine equilibrium distribution of chlorocarbons in such a model. 

Although a 3-D model that possesses adequate chemistry for the 

chlorocarbon/ozone problem would contain some of the dynamical limitations previously outlined and may not predict the global mean-temperature 

profile, it would provide an assessment of the importance of chemical/ 

dynamical feedbacks and of the inhomogeneity of the atmosphere for this 

problem. Provided that more detailed dynamical models continue to be 

developed, even more realistic models of the atmosphere can be expected 

to be available to answer questions regarding potential climatic changes of 

the stratosphere during the next decade. Present 34) models place much 

more emphasis on dynamics than on radiation and chemistry. Future 

efforts should aim at a more balanced treatment. 

TRANSPORT PARAMETERIZATIONS IN 3-D MODELING 

OF THE STRATOSPHERE 

Chapter 4 discusses the parameterization of transport in I -D models by the 

height-dependent eddy-diffusion parameter, K. A previous section in this 

chapter discusses the use of eddy-diffusion coefficients in 2-D models. It is 

not generally appreciated, however, that some parameterizations of transport will also probably be needed in 3-D models of the stratosphere. Such 

parameterizations are required because of the sizable spatial resolution 

(hundreds of kilometers in the horizontal and several kilometers in the 

vertical) required for fitting such a problem into even the largest digital 
MULTIDIMENSIONAL CFM-0 219 3 MODELING 

computers, as, well as for accomplishing a simulation in a reasonable 

amount of time. This situation is well-known in the meteorology of the 

troposphere, in which the spatial resolution is inadequate to explicitly treat 

atmospheric processes such as clouds, fronts, squall lines, and tornadoes. 

In 3-D stratospheric modeling, there are two basic types of transport parameterization problems. The first is the representation of mixing due to 

motions with space scales less than the minimum scale that can be explicitly resolved by the model. The second type is the parameterization of 

vertical transport across the tropopause. 

SUBGRID-SCALE MIXING 

In the numerical solution of any atmospheric transport problem, the distinction between nonturbulent and turbulent transport is arbitrary. It 

depends on the grid size and time step (i.e., on the averaging volume in 

four-dimensional space). Transport by all waves whose length isless than 

two grid distances must be described statistically as turbulent transport. 

As the scale of the averaging volume is increased, the statistical description 

of the turbulent transport is extended to include the effects oflarger and 

larger scale waves. The turbulence then changes by becoming larger in 

magnitude (large velocity components and displacement) and by becoming 

more anisotropic and inhomogeneous. Generally speaking, the major contribution to turbulent transport is from the largest scales of motion that 

cannot be resolved by the grid system (i.e., waves two to three times grid 

distance). For most general-circulation models with grid distances of more 

than 300 kilometers, the "turbulence" is still strongly anisotropic, and it 

can be questioned whether the turbulent transport is adequately described. 

However, above 15 kilometers, the transport in the stratosphere is increasingly dominated by planetary scale motions that are explicitly resolved in 

the 3-D models. Thus, except in the lower stratosphere, mixing by quasihorizontal subgrid scale motion may not be very important. 

In addition to the subgrid-scale "waves" that must be treated statistically, 

there are small-scale turbulent eddies (clear air turbulence) that may also 

play a role in vertical mixing processes. However, stratospheric turbulence 

observations (Lilly et al., 1975; and Heck and Panofsky, 1975) indicate 

that average vertical mixing coefficients for heat in the stratosphere below 

20 kilometers are of order 102cm2 see 1 , at least an order of magnitude 

less than mixing coefficients due to all scales of motions. 
220 CHLOROFLUOROMETHANE WORKSHOP REPORT 

If the reasonable assumption is made that mixing coefficients for heat are 

of the same order of magnitude as mixing coefficients for contaminants, 

the conclusion is reached that small-scale turbulence is unimportant for 

vertical transport of trace gases below 20 kilometers. No analogous direct 

estimates exist above 20 kilometers. 

VERTICAL TRANSPORT ACROSS THE TROPOPAUSE 

Because many stratospheric trace species have their sources in the troposphere, it is essential to properly represent tropospheric/statospheric exchange processes. Reiter (1975) has estimated that 70 percent of the total 

mass of the stratosphere is exchanged with the troposphere each year. 

Several mechanisms appear to be responsible for the irreversible transport 

of trace species between the stratosphere and troposphere. The most 

spectacular types of motion that can produce stratospheric/tropospheric 

exchange are the intrusions of stratospheric air into the lower troposphere 

associated with growing extratropical cyclones. These intrusions, which 

are found in the region of upper-level fronts associated with the tropospheric jet stream, occur in thin laminae. (Typical dimensions are 1 kilometer in the vertical, more than 100 kilometers in the horizontal orthogonal to the wind, and more than 1000 kilometers parallel to the wind). 

These laminae are then destroyed by vertical mixing on the culmulus scale 

or by mixing into the surface boundary layer within 1 to 3 days. A large 

fraction of the stratospheric air is removed by this process in 1 year, but 

the scales that produce the transport are below or close to the limit of the 

resolution for a general-circulation model. If this type of transport is 

underestimated, it could be expected that high-latitude exchange of the 

stratospheric constituents (such as ozone) would be underestimated and 

the low-latitude exchange at the confluence of the Ferrel and Hadley cells 

would be overestimated. Such a pattern of errors does occur with lowresolution general-circulation models. Decreasing averaging volume should 

decrease the errors and the complexity required in the statistical description, but the burden is then transferred to the computational domain. 

A related phenomenon is the transport produced by the mean circulation 

associated with various jet-stream and longwave systems. A systematic 

transverse circulation is associated with these jets that produces mean 

rising motion on the anticyclonic .shear side of the jet stream and mean 

sinking motion on the cyclonic side (Mahlman, 1975). This mechanism is 

not only capable of returning tropospheric air to the stratosphere, but 

also appears to be responsible for much of the upward energy flux that 

excites planetary waves in the stratosphere. Because this process occurs on 
MULTIDIMENSIONAL CFM-0 221 3 MODELING 

large scales, the resulting transport can be explicitly incorporated in 3-D 

models. 

Cumulonimbus convection has often been proposed as a means of transporting trace species from the troposphere to the stratosphere. However, 

there is some evidence from laboratory convection studies and radioactivity deposition observations (Reiter and Mahlman, 1965) that the collapse of cumulus cells in the stratosphere due to. negative buoyancy and 

the subsequent sinking back into the troposphere may in fact cause 

entrained stratosphere air to enter the troposphere. So far, there is little 

evidence to indicate that cumulus convection provides an effective direct 

source of upward transport across the tropopause. However, convective 

clouds clearly transport trace species from the lower troposphere to the 

vicinity of the tropopause. 

In the tropical regions, substances that have been brought to the upper 

troposphere by cumulus convection may be transported upward through 

the lower stratosphere by the slow mean vertical motion associated with 

the rising branch of the Hadley cell. This slow rising motion through the 

cold tropical tropopause has long been accepted as a likely explanation for 

the observed low concentration of water vapor in the stratosphere (that is, 

water-vapor mixing ratios in the stratosphere do not exceed the saturation 

value at the tropical tropopause.) Multidimensional general-circulation 

models that explicitly compute the tropospheric circulation can explicitly 

incorporate the large-scale Hadley circulation. However, in models in 

which the tropopause is used as the lower boundary, the vertical component of the Hadley circulation must be included as a boundary condition. 

In either case, vertical transports by cumulus clouds can only be included 

by suitable parameterizations. 

OBSERVATIONAL NEEDS OF MULTIDIMENSIONAL MODELS 

OF THE STRATOSPHERE 

The specific concern of this subsection is the observational basis required 

for ensuring that multidimensional models are producing reasonable information about the overall impact of CFM's. It is important to note at the 

outset that theoretical modeling and observational studies are, and should 

be, interactive ventures. Not only are quantitative checks on model outputs required, but also measurement strategies that address important 

theoretical problems must be designed and completed. The great variability in space and time that exists in the stratosphere in both the meteorological fields (wind, temperature, and pressure) and the ozone distribution 
222 CHLOROFLUOROhMETHANE WORKSHOP REPORT 

must be understood. It is suspected that most of the ozone variability, 

at least below 30 kilometers, is due to transport by large-scale motions. 

Even in the region of local photochemical control, advection of photochemically active minor species such as NOr and Ct will influence the 

ozone distribution. 

Most 3-D models will produce both wind and temperature fields systematically before utilizing these fields in photochemical transport models. 

Quasi-geostrophic theory can provide such models in the midlatitudes, and 

balance equations have been used to extend these models to the entire 

sphere (Cunnold et al., 1975). However, primitive equation models are 

probably needed for simulating the tropical stratosphere. To compare 

model dynamics with observations will require temperature measurements 

with grid resblutions on the order of 1000, 250, and 3 kilometers in the 

zonal, meridional, and vertical directions, respectively. 

Time resolutions on the order of 1 day are also needed. Diagnostic studies 

utilizing these data will isolate various dynamical and energetic mechanisms, thus increasing the understanding of the stratosphere. When compared to similar studies of multidimensional model output, these studies 

can help to verify various aspects of the models. Horizontal velocities 

can be determined from the temperature field and reference-level winds by 

approximate methods whose accuracy can be matched to that of the 

temperature observations. Vertical velocities can also be obtained by a 

number of indirect methods. When combined with global observations of 

chemical constituents, these velocities can be used to compute transports. 

Direct wind meaiurements would provide a valuable check on model output and would verify indirectly determined winds. Horizontal wind components can be measured using the standard meteorological observational 

network augmented by data utilizing radar, superpressure balloons, and 

possibly satellite-based Doppler techniques. Such measurements would be 

most useful in the tropics, but there they would have to be accurate within 

- 1 . 2 to 4 nisec 

In order to establish that transport variability is being reasonably simulated by any given model, global coverage of ozone and of at least one 

photochemically simple trace constituent with the spatial and temporal 

resolution previously suggested for temperature will be needed. Remote 

sensing from satellites is clearly needed for establishing ths type of data 

base. Ozone has been already observed from a number of satellites, albeit 

with poor vertical resolution below the ozone peak N2 0 and CH4 are the 

traditional candidates for a photochemically important tracer. N20 is 

preferred primarily because methane can be lost by reaction with species 
MULTIDIMENSIONAL CFM-O 223 3 MODELING 

such as OH, as well as by reaction with O('1D) and through photolysis, 

both of which are also sinks for N20. The important loss mechanism for 

N20 is probably photolysis; if so, this greatly simplifies its transport equation. At present concentrations, the CFM's themselves, particularly CF 2CI2 

and CFC13, may be useful tracers in the stratosphere because they appear 

to be inert except for photolysis. Furthermore, the distribution of their 

sources is well-known. Of course, ozone itself should be a good tracer in 

the lower stratosphere, and its distribution there is the final test of any 

model. 

The Stratospheric and Mesospheric Sounder (SAMS) to be launched in 

1978 on Nimbus-G will measure N20, as well as other trace constituents, 

but better vertical resolution will be required. The Limb Infrared Monitor 

of the Stratosphere (LIMS), also to be flown on Nimbus-G, has better 

height resolution, but it measures NO2 and HNO rather than N O. The 

former are important in the scheme for ozone destruction, but their complex photochemistry limits their use as tracers for advective transport. 

The data processing task -associated with the global coverage described 

previously is enormous. Ideally, those responsible for data analysis should 

work in close association with the experimenters. This would permit the 

data to be reduced in a usable format familiar to the researchers and would 

alert them to possible problems or limitations inherent in the instrumentation or its operation. The data must not simply be archived. Again, any 

data analysis should place some emphasis on determining the scales of 

spatial and temporal variability. It should be noted that significant data 

analysis has been and isbeing performed on presently existing data (Angell 

and Korshover, 1976; Hartmann, 1976; Quiroz, 1975; and Barnett et al., 

1975). 

The degree to which any model can simulate mass transport of photochemically active trace constituents, including ozone, depends not only on 

accurately describing advection and/or diffusion processes, but also on 

quantitatively specifying the sources and sinks of the trace constituents. 

For this reason, multidimensional models must eventually include both 

adequate dynamic transport and accurate photochemistry. 

One class of observations that is obviously needed for any model attempting to assess the impact of CFM is a series of measurements to validate the 

photochemistry itself. To do this, it is necessary to measure families of 

photochemical species such as CFM or NO., or representatives of each of 

the odd nitrogen, chlorine, hydrogen, and oxygen families simultaneously. 
224 CHLOROFLUOROMETHANE WORKSHOP REPORT 

In contrast to the global coverage required for describing advective transport (even when parameterized as eddy diffusion), observations designed 

to test understanding, of the photochemistry are local in nature and are 

best done from balloon, airplane, or space-shuttle platforms. 

Microwave measurements from the ground may also help. To minimize 

meteorological influences on these observations, the minor species are best 

sampled in the region of local photochemical equilibrium (above 25 to 30 

kilometers). Temperature should be measured simultaneously for computing kinetic reaction rates. Photodissociation rates can be accurately 

specified if the ozone overburden is also measured. However, the solar 

flux and its variability should be measured more precisely at wavelengths 

less than approximately 200 nanometers where both CFM and N20 are 

photodissociated. Some stratospheric circulation models have bypassed 

generating the tropospheric circulation, with its vastly varying scales of 

motion and complex hydrological cycle, by placing their lower boundary 

in the upper troposphere. These models are advantageous in that they permit highly useful mechanistic studies, but the suppression of a proper 

troposphere/statosphere interaction limits their usefulness. If such 

models are to describe the more general transport in the stratosphere, they 

will eventually require specification of the variable temperatures and winds 

along a meteorological reference level in the upper troposphere. The distributions of CFM, N20 and H.0, all of which are precursors of photo­ , 

chemically active minor species, would also have to be specified along such 

a boundary. 

The observations required for describing such a boundary should be part 

of a program directed toward understanding transport across the tropopause. A great deal more must be known about how the precursors and 

reservoirs of odd nitrogen, hydrogen, and chlorine enter or leave the lower 

stratosphere. In particular, the roles of cumulus convection, jet streams, 

and midlatitude barocinic systems need more study. (See the previous 

section on "Vertical Transport Acrosi the Tropopause.") This suggests 

latitudinal surveys of CFM, N20, HNO3 , and HCl near the tropopause and 

perhaps correlation of these observations with synoptic-scale ineteorological events (e.g., storm systems and temporally varying jet streams). Aircraft and balloons provide suitable instrument platforms. 

The standard meteorological observational network and daily analysis can 

provide some of the necessary information. In particular, it outlines the 

temporal variability in the upper troposphere of the otherwise lowfrequency, vertically propagating planetary waves believed to produce 

much of the dynamic activity in the upper atmosphere. 
MULTIDIMENSIONAL CFM-O 3 MOD;ELING 225 

Most of the foregoing discussion has addressed models that explicitly 

generate the large-scale wind field used to study mass transport. Observations of at least two -trace constituents are needed for models that parameterize their advective transport as an eddy-diffusion process: one for 

parameterization and one for independent comparison. Measurements 

emphasizing the latitudinal and seasonal variations of these tracers and 

other important photochemically active species are needed in sufficient 

number that reasonable averages can be computed for comparison with 

model calculations. Computations for a strongly perturbed stratosphere, 

however, require understanding of how the coefficients that describe the 

effective transport in these parameterized models will evolve. Variability 

with time is not easily included in these models. 

In summary, the following measurements are emphasized, in addition to 

those needed to-verify the stratospheric photochemistry itself: 

Global coverage of temperature, ozone, and at least two photochemically simple tracers, such as N.O, with adequate resolution for 

transport studies (namely, observational spacings of at least 1000, 

250, and 3 kilometers in the zonal, meridional, and vertical directions, respectively) 

Specification of temperature and wind along at least two meteorological reference levels in the upper troposphere; observations of 

CFM, NOY, N20, H20 , HNO3, and HCI at this reference level and 

near the tropopause. These observations should be part of a more 

general program directed toward understanding transport into and 

out of the lower stratosphere. (See the section on "Vertical Transport Across the Tropopause.") 

Several years of coverage are needed for both of these items. Furthermore, 

analysis should examine spatial and temporal variability. Measurement 

strategies must include planning and support for this data analysis. 

This observational program is an extensive one, but should be undertaken. 

As a more immediate step toward establishing a useful data base for ongoing modeling efforts, however, it is recommended that one or more 

International Stratospheric Periods (ISP) be established during which coordinated measurements would be made. The measurement period suitable 

to averaging is limited by natural variability to I 'to 3 days during winter, 

1 week during early fall and late spring, and 2 weeks during summer. The 

aim of such a measurement program would be to obtain vertical profiles 
226 CHLOROFLUOROMETHANE WORKSHOP REPORT 

of temperature and as many constituents as possible for at least one highlatitude, one midlatitude, and one low-latitude station. By also utilizing 

simultaneous aircraft and satellite observations, it should be possible to 

obtain an instantaneous global ,picture of the stratosphere at the time of 

the ISP. 

CONCLUSIONS 

ASSESSMENT UNCERTAINTY DUE TO TRANSPORT 

One-dimensional models utilizing detailed chemistry with simplified radiation treatments and with all effects of dynamics parameterized by means 

of a height-dependent eddy diffusion have been used to estimate the CFM 

effects on ozone in this report, as well as in the NAS Report, Halocarbons: 

Effects on Stratosphere Ozone (1976). The results of these models provide the best available estimate of the CFM effects on ozone that depends 

on the anticipated CFM-release scenario. These model results also provide 

the uncertainty of this best estimate. This uncertainty is a result not only 

of possible errors in the parameterized transport estimates, but also of 

uncertainties in CFM release rates, stratosphere chemistry, and in the 

approximation of the real atmosphere by a relatively simple 1-D model. 

The NAS Report (1976) has quoted the 95-percent uncertainty limits due 

to the use of a I-D model for approximating the transport and distribution 

of chemical species to be given by the best estimate multiplied by a factor of (.7)"'. This uncertainty is for the problem of a continuous CFMrelease scenario at constant 1973 levels. The uncertainty with an interrupted release scenario would be less than this. This is due to the nature 

of the CFM problem in which the CFM's are expected to become uniformly mixed throughout the troposphere before being transported in 

substantial amounts through the region of minimum transport in the vicinity of the tropopause. This assumes that CFM's are lost from the troposphere mainly by upward transport rather than by in situ removal. Because of the expected uniform mixing of CFM's in the troposphere, the 

uncertainties in vertical transport resulting from the use of a 1-D model are 

relatively small. 

The uncertainties resulting from the use of the l-D model quoted in the 

NAS Report (1976) were derived by extensive experimentation with different eddy-diffusion profiles. It is not entirely convincing that such a 

procedure gives the total uncertainty that results from approximating a 

3-D problem by a I-D model. For instance, the use of a l-D model 

assumes the average of a function of several variables to be equal to the 
MULTIDIMENSIONAL CFM-O3 MODELING 227 

same function of the averaged variables. This approximation, while probably not leading to great errors, should increase the uncertainties from 

that quoted in the NAS Report (1976). Uncertainty about whether or not 

transport can be handled correctly in a multiconstituent 1-D model using 

a single eddy-diffusion profile also leads to a widening of the uncertainty 

interval. Indeed, the 3-) tracer studies of Mahlman (1975) show that 

derived 1-D eddy-diffusion coefficients can differ significantly among 

various tracers. Furthermore, there are a number of feedback effects that, 

in some cases, are very difficult to quantitatively evaluate and that cannot 

be treated in a 1-D model. The ultimate feedback effects are the direct 

and indirect CFM effects on climate that are beyond the scope of this 

report. The neglect of these feedback effects also leads to a widening of 

the NAS unceitainty interval that results from using l-D models to 

approximate the atmosphere. The extent of the widening of the uncertainty limits implied by the neglect of feedback processes is an open question. 

CONTRIBUTIONS OF MULTI DIMENSIONAL MODELS 

There is no expectation for the development of a comprehensive multidimensional model incorporating all of the necessary physics and chemistry required for assessment applications, Assessment's will always be 

based on a judicious complement of studies with different models, including l-D, 2-D, and 3-D models. This procedure permits different physical 

processes and effects to be treated in adequate detail and provides results 

that can be incorporated in other models, emphasizing other aspects of the 

overall problem. It is essential that special care be taken to ensure that all 

important feedbacks and interactions are included in this process. Although it will probably require at least several years to obtain satisfactory 

investigation of the comprehensive radiative-dynamical/chemical problems, 

it is important to emphasize that contributions can be expected in the 

near term from 2-D and 3-D models, as well as from 1-D models. 

Two-dimensional models permit the inclusion of latitudinal and seasonal 

dependencies of transport, temperature, and the solar flux. Because of the 

observed variation of 0. concentration with latitude, the strong change of 

solar flux and dynamical transports with latitude, and the dependence of 

reaction rates on temperature, there is reason to expect that realistic 2-D 

models can provide better assessments or useful information for improving 

assessments made with 1-]1 models. The 2-D models must rely on pa-ameterized transport, as isthe case with 1-1 models; the use of four transport 
228 CHLOROFLUOROMETHANE WORKSHOP REPORT 

coefficients instead of the one in a l-D model means that more observational data are required for parameterization, but eventually a better determination of transport effects on 0 assessments should result. The 2-D 

models have a number of limitations arising from the parameterization of 

effects due to longitudinal variations. Primarily because of such parameterizations, the radiative, dynamical, and chemical 2-D models are not as 

"fully coupled" as desired. 

Three-dimensional models have the advantage of minimizing the parameterization of transport. The scales of motion responsible for most of the 

transport in the stratosphere can, in fact, be explicitly resolved with feasible grid spacings. This capability will permit a maximum incorporation of 

feedbacks and interactions that will provide information for upgrading the 

reliability of 1-D and 2-D models. The 3-D models have already been of 

some value for diagnostic purposes (for example, for tracer studies, for 

providing an indication oflongitudinal variabilities, and for testing assumptions made in simpler models), and the applications can be expected to 

increase as the realism of the 3-D models improves. The primary limitations of the 3-D models are related to the extensive computer requirements, which can translate into a great expense if the computations are 

made for extended simulations and into a requirement for inclusion of a 

simplified chemistry. As previously emphasized, a proper strategy for 

investigation must involve a careful complement of l-D, 2-D), and 3-D 

studies. 

OBSERVATIONAL NEEDS 

Observations needed by multidimensional models can be separated into 

the following categories: (I) those needed for verification of the photochemisty itself, (2) those needed for parameterization of dynamical 

processes such as eddy transport in 2-) models and subgrid-scale processes in 3-) models, and (3) those needed for determining larger scale winds, 

whether to derive mean motions for 2-D models or to examine the scales 

of temporal and spatial variability for comparison with 3-D model output. 

Each of these categories requires different observational coverage and 

resolution. Studies for supporting 3-D models require global coverage of 

temperature and at least two tracers, such as N2O , with spatial resolutions of 1000, 250, and 3 kilometers in the zonal, meridional, and vertical directions, respectively. Even 2-D models need measurements of the 

longitudinal variability of tracer concentrations in order to assess the 

representativeness of a given zonal average, as well as to evaluate the 

validity of various eddy-transport parameterizations. 
MULTIDIMENSIONAL CFM-0 3 MODELING 229 

Observations required to specify lower boundary conditions for models 

that do not explicitly include the troposphere should be part of a more 

general program to study transport across the tropopause and in the lower 

stratosphere. In particular, more knowledge is needed concerning the 

ioles played by tropical cumulus convection, jet streams, and midlatitude 

baroclinic systems. 

In all cases, data analysis should examine spatial and temporal variability. 

Furthermore, the difficulty of processing the large quantity of data required by multidimensional models, particularly 3-1) models, makes it 

imperative that measurement strategies include sufficient planning and 

support for data distribution and analysis. 
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APPENDIX
 

CHLOROFLUOROMETHANE WORKSHOP
 

ATTENDEES
 

LABORATORY MEASUREMENTS 

W. DeMore - Chairman
 

Jet Propulsion Laboratory
 

L Stief - Vice-Chairman
 

Goddard Space Flight Center
 

R. Watson - Halogen Reaction Rates
 

Jet Propulsion Laboratory
 

R. Hampson - NO, Reaction Rates
 

National Bureau of Standards
 

D.Garvin
 

National Bureau of Standards
 

J.Margitan - HOx Reaction Rates
 

University of Michigan
 

M.Molina - Photolysis Rates and Quantum Yields 

University of California, Irvine 

D. Golden
 

Stanford Research Institute
 

OZONE MEASUREMENTS AND TRENDS 

J, London - Chairman 

University of Colorado 

C. Reber - Vice-Chairman
 

Goddard Space Flight Center
 

J. Glle - Ozone Measurements from Satellites 

National Center for Atmospheric Research 

261 

Precding page blank
 
262 CHLOROFLUOROMETHANE WORKSHOP REPORT 

W. Komhyr - Ground-Based Ozone Measurements 

National Oceanic and Atmospheric Administration 

A. Krueger - Ozone Profile Measurements 

Goddard Space Flight Center/Colorado State University 

J. Angell - Ozone Trend Analyses 

National Oceanic and Atmospheric Administration 

B.Conrath
 

Goddard Space Flight Center
 

A. Belmont
 

Control Data Corporation
 

J. Campbell
 

Langley Research Center
 

MINOR SPECIES AND AEROSOL MEASUREMENTS 

R. Megill - Chairman
 

Utah State University
 

G. Newton - Vice-Chairman
 

Goddard Space Flight Center
 

A. Lazrus - Halocarbons
 

National Center for Atmospheric Research
 

A. Schmeltekopf - N2 0, CH 4 , and others 

National Oceanic and Atmospheric Administration 

J. Anderson - C10 X Measurements
 

Michigan University
 

B. Ridley - NOx Measurements
 

York University (Canada)
 

W. Heaps - HOx Measurements
 

Goddard Space Flight Center
 

'v+',t,.TU-tre+, 4tpt++.--;+ 
ATTENDEES 263 

W.Moore - 0/0 Ratio
 

Utah State University
 

J. Rosen -Aerosols
 

University of Wyoming
 

E. Good - Winds and Dynamics
 

Air Force Geophysics laboratory
 

D. Heath - Solar-Flux Measurements
 

Goddard Space Flight Center
 

R. Rundel
 

Johnson Space Center
 

W.Evans
 

Atmospheric Environment Service (Canada)
 

D. Muroray
 

University of Denver
 

ONE-DIMENSIONAL MODEL CALCULATIONS 

R. Stolarski - Chairman
 

Goddard Space Flight Center
 

J. Herman - Vice-Chairman
 

Goddard Space Flight Center
 

J. Chang - Transport Coefficients
 

Lawrence Livermore laboratories
 

F. Luther - Photodissociation Rate Calculations
 

Lawrence Livermore Laboratories
 

S. Liu - Diurnal Effects
 

University of Michigan
 

V. Ramanathan - Radiative/Convective Feedbacks
 

National Center for Atmospheric Research
 

C.Chou -Tropospheric and Stratospheric Sinks
 

University of California, Irvine
 
264 CILOROFLUOROMETHANE WORKSHOP REPORT 

R. Cicerone - Comparison Theory and Experiment for CIO x 

University of Michigan 

R. Kurzeja - Comparison Theory and Experiment for NOx National Center for Atmospheric Research 

R. Turco - Comparison Theory and Experiment for HOx 

R&D Associates 

D. Butler - Uncertainty Analyses
 

Goddard Space Flight Center
 

L. Callis
 

Langley Research Center
 

R. Boughner
 

Langley Research Center
 

J. Nealy
 

Langley Research Center
 

W.Duewer
 

Lawrence Livermore Laboratories
 

J. McConnell
 

York University (Canada)
 

P.Jesson 

E. L Du Pont de Nemours 

R. Hirsch 

E. I. Du Pont de Nemours 

C.Miller 

E. I. Du Pont de Nemours 

L. Glasgow,
 

EI.lDu Pont do Nemours
 

S.Wofsy
 

Harvard University
 
ATTENDEES 265 

D.Sze 

Environmental Research and Technology, Inc. 

J. McAfee
 

National Oceanic and Atmospheric Administration
 

S.Chandra
 

Goddard Space Flight Center
 

MULTIDIMENSIONAL MODEL CALCULATIONS 

M. Geller - Chairman
 

University of Illinois
 

i. Hansen - Vice-Chairman
 

GSFC/Goddard Institute for Space Studies
 

G. Widhopf- Status of 1-D Models
 

Aerospace Corporation
 

D. Cunnold - Status of 3-1) Models
 

Massachusetts Institute of Technology
 

H- Panofsky - Parameterization of Transport Processes
 

Pennslyvania State University
 

G. Mahlnan - Radiation, Chemistry, and EDynanics
 

National Oceanic and Atmospheric Administration
 

R. Zurek - Observational Needs
 

Jet Propulsion Laboratory
 

G. Brasseur
 

Institut d'Aeronomie Spatiale (Belgium)
 

J. Holton
 

University of Washington
 

D. Hartman
 

National Center for Atmospheric Research
 

C-Reigel
 

San Jose State University
 
266 CHLOROFLUOROMETHANE WORKSHOP REPORT 

R. Whitten
 

Ames Research Center
 

GODDARD SPACE FLIGHT CENTER EDITORIAL COMMITTEE 

R. Hudson - Chairman 

N. Spencer 

W.Bandeen 

R. Stolarski 

J. Herman 

C.Reber 

G.Newton 

L. Stief 

I. Hansen
 

Ex Officio
 

R. Greenwood, NASA Headquarters 

R. Schiffer, NASA Headquarters 
National AerdflaiitiGSifd 

Space Administration 

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