World Meteorological Organization

Global Ozone Research and Monitoring Project—Report No. 47

SCIENTIFIC ASSESSMENT OF

OZONE DEPLETION: 2002

Pursuant to Article 6 of the Montreal Protocol

on Substances that Deplete the Ozone Layer

National Oceanic and Atmospheric Administration

National Aeronautics and Space Administration

United Nations Environment Programme

World Meteorological Organization

European Commission
iii

Alberto Adriani Italy

Ayité-Lô Nohende Ajavon Togo

Daniel L. Albritton US

Douglas R. Allen US

Georgios T. Amanatidis Belgium

Stephen O. Andersen US

James Anderson US

Gustavo A. Argüello Argentina

Antti Arola Finland

Roger Atkinson US

Pieter J. Aucamp South Africa

John Austin UK

Linnea M. Avallone US

Alkiviadis F. Bais Greece

Mark P. Baldwin US

Stephen R. Beagley Canada

Pranvera Bekteshi Albania

Germar Bernhard US

Donald R. Blake US

Nicola J. Blake US

Mario Blumthaler Austria

Greg E. Bodeker New Zealand

Rumen D. Bojkov Germany

Olivier Boucher France

Michel Bourqui UK

Geir O. Braathen Norway

Peter Braesicke UK

Guy P. Brasseur Germany

Bram Bregman The Netherlands

Christoph Brühl Germany

William Brune US

James H. Burkholder US

Neal Butchart UK

James H. Butler US

James Calm US

Pablo O. Canziani Argentina

Ken S. Carslaw UK

Marie-Lise Chanin France

Mian Chin US

Martyn P. Chipperfield UK

John Christy US

Peter S. Connell US

Brian Connor New Zealand

David B. Considine US

R. Anthony Cox UK

Derek M. Cunnold US

Martin Dameris Germany

John S. Daniel US

Richard G. Derwent UK

Susana B. Diaz Argentina

LIST OF INTERNATIONAL AUTHORS,

CONTRIBUTORS, AND REVIEWERS

Assessment Cochairs

Ayité-Lô Nohende Ajavon

Daniel L. Albritton

Gérard Mégie

Robert T. Watson

Chapters and Lead Authors

Chapter 1. Controlled Substances and Other Source Gases (Stephen A. Montzka and Paul J. Fraser)

Chapter 2. Very Short-Lived Halogen and Sulfur Substances (Malcolm K.W. Ko and Gilles Poulet)

Chapter 3. Polar Stratospheric Ozone: Past and Future (Paul A. Newman and John A. Pyle)

Chapter 4. Global Ozone: Past and Future (Martyn P. Chipperfield and William J. Randel)

Chapter 5. Surface Ultraviolet Radiation: Past and Future (James B. Kerr and Gunther Seckmeyer)

Twenty Questions and Answers About the Ozone Layer (David W. Fahey)

Coordinating Editor

Christine A. Ennis

Authors, Contributors, and Reviewers
iv

AUTHORS, CONTRIBUTORS, AND REVIEWERS

Ed Dlugokencky US

Anne R. Douglass US

Katja Drdla US

Thomas Duafala US

Ellsworth S. Dutton US

Victor L. Dvortsov US

Ezequiel Echer Brazil

Stephen Eckermann US

Kalju Eerme Estonia

James W. Elkins US

Andreas Engel Germany

Christine A. Ennis US

David W. Fahey US

Joe Farman UK

Vitali E. Fioletov Canada

Eric L. Fleming US

Ian Folkins Canada

Piers M. de F. Forster UK

Paul J. Fraser Australia

Randall R. Friedl US

Ru-Shan Gao US

Rolando R. Garcia US

Marvin A. Geller US

Melvyn Gelman US

Christian George France

Sophie Godin-Beekmann France

Marco González Kenya

Florence Goutail France

Michael Graber Kenya

Hans-F. Graf Germany

Volker Grewe Germany

Joanna D. Haigh UK

Patrick Hamill US

David B. Harper UK

Neil R.P. Harris UK

Didier Hauglustaine France

Peter H. Haynes UK

Jay R. Herman US

David J. Hofmann US

James R. Holton US

Robert D. Hudson US

Drusilla Hufford US

Abdelmoneim A. Ibrahim Egypt

Mohammad Ilyas Malaysia

Takashi Imamura Japan

Ivar S.A. Isaksen Norway

Charles H. Jackman US

Daniel J. Jacob US

Mauricio Jaramillo-Ayerbe Colombia

Paul Johnston New Zealand

David Karoly Australia

Nozomi Kawamoto Japan

Jack A. Kaye US

James B. Kerr Canada

M.A.K. Khalil US

Peter W. Kiedron US

Dieter Kley Germany

Bjørn Knudsen Denmark

Malcolm K.W. Ko US

Yutaka Kondo Japan

Yuri Koshelkov Russia

Karin Kreher New Zealand

Nickolay A. Krotkov US

Janusz W. Krzys ́cin Poland

Lambert Kuijpers The Netherlands

Michael J. Kurylo US

Karin Labitzke Germany

Murari Lal India

Shyam Lal India

Ulrike Langematz Germany

Kathleen O. Lantz US

Neils Larsen Denmark

Katherine S. Law UK

Mark G. Lawrence Germany

J. Ben Liley New Zealand

Roger Lin US

Shaw Liu Taiwan R.O.C.

Jennifer A. Logan US

Craig S. Long US

Daniel Lubin US

A. Robert MacKenzie UK

Sasha Madronich US

Emmanuel Mahieu Belgium

Gloria L. Manney US

Elisa Manzini Germany

Céline Mari France

Timothy J. Martin Austria

W. Andrew Matthews New Zealand

Konrad Mauersberger Germany

Archie McCulloch UK

Gordon McFadyen UK

Mack McFarland US

Daniel S. McKenna US

Richard L. McKenzie New Zealand

Richard D. McPeters US

Ralf Meerkötter Germany

Gérard Mégie France

Inna A. Megretskaia US

Davit Melkonyan Armenia

Abdelwahid Mellouki France

Pauline M. Midgley Germany

Alvin J. Miller US

Stephen A. Montzka US

Rolf Müller Germany

Nzioka John Muthama Kenya

Tatsuya Nagashima Japan

Hideaki Nakane Japan

Eric R. Nash US
v

AUTHORS, CONTRIBUTORS, AND REVIEWERS

John Nash UK

Patrick J. Neale US

Paul A. Newman US

Samuel J. Oltmans US

Alan O’Neill UK

Michael Oppenheimer US

David E. Oram UK

Eduardo Palenque Bolivia

Panos Papagiannakopuolos Greece

David Parker UK

Steven Pawson US

Stuart A. Penkett UK

Sunil Kumar Peshin India

Thomas Peter Switzerland

Klaus Pfeilsticker Germany

Giovanni Pitari Italy

Ulrich Platt Germany

Ian Plumb Australia

Jean-Pierre Pommereau France

Lamont R. Poole US

Robert W. Portmann US

Gilles Poulet France

Michael J. Prather US

Margarita Préndez Chile

Ronald Prinn US

Michael H. Proffitt Switzerland

John A. Pyle UK

S. Ramachandran India

V. Ramaswamy US

William J. Randel US

Lakshman Randeniya Australia

Philip J. Rasch US

A.R. Ravishankara US

Claire E. Reeves UK

Markus Rex Germany

Brian A. Ridley US

Curtis P. Rinsland US

Henning Rodhe Sweden

José M. Rodríguez US

Bjørg Rognerud Norway

Joan E. Rosenfield US

Martin N. Ross US

Eugene Rozanov Switzerland

James M. Russell III US

Nelson A. Sabogal Kenya

Ali A. Sabzipavar Iran

Ross J. Salawitch US

Eugenio Sanhueza Venezuela

Michelle L. Santee US

Toru Sasaki Japan

Yasuhiro Sasano Japan

Sue M. Schauffler US

Ulrich Schmidt Germany

Christina Schnadt Germany

Ulrich Schumann Germany

M. Daniel Schwarzkopf US

Paul W. Seakins UK

Gunther Seckmeyer Germany

Dian J. Seidel US

Dudley Shallcross UK

Theodore G. Shepherd Canada

Drew T. Shindell US

Keith P. Shine UK

Masanori Shitamichi Japan

Peter G. Simmonds UK

Paul C. Simon Belgium

Harry Slaper The Netherlands

James R. Slusser US

Claire A. Smith UK

Sergei Smyshlyaev Russia

Susan Solomon US

Johannes Stähelin Switzerland

L. Paul Steele Australia

Andreas Stohl Germany

Richard S. Stolarski US

Susan E. Strahan US

B.H. Subbaraya India

Christine C. Sweet US

Petteri Taalas Finland

Azadeh Tabazadeh US

David W. Tarasick Canada

Gary Taylor Canada

Manfred Tevini Denmark

Christiane Textor Germany

Larry W. Thomason US

David Thompson US

Claudia Timmreck Germany

Darin W. Toohey US

Geoffrey C. Toon US

Adrian F. Tuck US

Jan C. van der Leun The Netherlands

Michel Van Roozendael Belgium

Michiel van Weele The Netherlands

Guus J.M. Velders The Netherlands

Jean Verdebout Italy

Mario Visca Italy

C. Michael Volk Germany

Peter von der Gathen Germany

Ray Wang US

Joe W. Waters US

Robert T. Watson US

Darryn W. Waugh US

Elizabeth C. Weatherhead US

Philipp Weihs Austria

Debra K. Weisenstein US

Ray F. Weiss US

Paul O. Wennberg US

Fei Wu US
vi

AUTHORS, CONTRIBUTORS, AND REVIEWERS

Donald J. Wuebbles US

Yoko Yokouchi Japan

Vladimir Yushkov Russia

Rodolphe J. Zander Belgium

Francesco Zaratti Bolivia

Joseph M. Zawodny US

Christos S. Zerefos Greece

Xiuji Zhou China

Tong Zhu China
vii

CONTENTS

SCIENTIFIC ASSESSMENT OF OZONE DEPLETION: 2002

PREFACE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ix

EXECUTIVE SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xiii

CHAPTER 1: CONTROLLED SUBSTANCES AND OTHER SOURCE GASES

Lead Authors: S.A. Montzka and P.J. Fraser

Scientific Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.1

1.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.5

1.2 Halogenated Ozone-Depleting Gases in the Atmosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.5

1.3 Halocarbon Sources Estimated from Industrial Production . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.19

1.4 Halocarbon Lifetimes, Ozone Depletion Potentials, and Global Warming Potentials . . . . . . . . . . . . . . . . . 1.21

1.5 Methyl Bromide and Methyl Chloride . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.36

1.6 Atmospheric Halocarbon Observations Compared with Expectations . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.44

1.7 Other Trace Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.51

1.8 Halogenated Source Gases in the Future . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.61

References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .1.71

CHAPTER 2: VERY SHORT-LIVED HALOGEN AND SULFUR SUBSTANCES

Lead Authors: M.K.W. Ko and G. Poulet

Scientific Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1

2.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.5

2.2 Dynamical and Chemical Characteristics of the Upper Troposphere (UT) and the Tropical

Tropopause Layer (TTL) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.7

2.3 Atmospheric Chemistry of Halogenated Very Short-Lived (VSL) Substances . . . . . . . . . . . . . . . . . . . . . . 2.15

2.4 Contribution of Halogenated VSL Substances to the Stratospheric Inorganic Halogen Budget . . . . . . . . . 2.25

2.5 Estimates for the Potential Impact of Halogenated VSL Substances on Column Ozone . . . . . . . . . . . . . . . 2.34

2.6 VSL Sulfur Species and Stratospheric Aerosols . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.39

References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.44

CHAPTER 3: POLAR STRATOSPHERIC OZONE: PAST AND FUTURE

Lead Authors: P.A. Newman and J.A. Pyle

Scientific Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.1

3.0 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.5

3.1 Trends of Ozone and Temperature in the Polar Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.6

3.2 Basic Polar Stratospheric Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.13

3.3 Quantification of Polar Ozone Loss: Observations and Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.41

3.4 Causes of Polar Stratospheric Temperature Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.58

3.5 Chemical-Climate Modeling of the Past and Future Polar Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.66

References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.81

Appendix 3A: Satellite Measurements in the Antarctic and Arctic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.104
CHAPTER 4: GLOBAL OZONE: PAST AND FUTURE

Lead Authors: M.P. Chipperfield and W.J. Randel

Scientific Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1

4.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.5

4.2 Description of Past Changes in Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.6

4.3 Current Understanding of Past Changes in Stratospheric Aerosol, Water Vapor, and NO2 . . . . . . . . . . . . . . 4.25

4.4 Current Understanding of Past Changes in Stratospheric Temperature . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.29

4.5 Chemical Influence on Past Changes in Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.37

4.6 Dynamical Influence on Past Changes in Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.52

4.7 Implications for Climate of Updated Stratospheric Ozone Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.62

4.8 Future Changes in Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.63

4.9 Synthesis of Current Understanding of Past and Future Changes in Ozone . . . . . . . . . . . . . . . . . . . . . . . . . 4.70

References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .4.72

Appendix 4A: Description of Ozone Datasets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.85

Appendix 4B: 2-D Model Scenarios . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.90

CHAPTER 5: SURFACE ULTRAVIOLET RADIATION: PAST AND FUTURE

Lead Authors: J.B. Kerr and G. Seckmeyer

Scientific Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.1

5.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.5

5.2 Current Understanding of Factors Affecting Surface UV Radiation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.5

5.3 Available Resources for Studying Surface UV Radiation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.11

5.4 UV Climatology, Temporal Changes, and Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.20

5.5 Expectations of UV in the Future . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.25

References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.31

Appendix 5A: Spectral Data Available from Databases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.43

Appendix 5B: Internet Addresses for UV Sites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.45

TWENTY QUESTIONS AND ANSWERS ABOUT THE OZONE LAYER

Lead Author: D.W. Fahey

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.1

I. Ozone in Our Atmosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.3

II. The Ozone Depletion Process . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.8

III. Stratospheric Ozone Depletion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.21

IV. Controlling Ozone-Depleting Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.28

V. Implications of Ozone Depletion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.32

VI. Stratospheric Ozone in the Future . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.36

APPENDICES

A LIST OF INTERNATIONAL AUTHORS, CONTRIBUTORS, AND REVIEWERS . . . . . . . . . . . . . . . . . . A.1

B MAJOR ACRONYMS AND ABBREVIATIONS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . B.1

C MAJOR CHEMICAL FORMULAE AND NOMENCLATURE FROM THIS ASSESSMENT . . . . . . . . . . C.1

viii

CONTENTS
ix

PREFACE

The present document will be part of the information upon which the Parties to the United Nations Montreal

Protocol will base their future decisions regarding protection of the stratospheric ozone layer.

The Charge to the Assessment Panels

Specifically, the Montreal Protocol on Substances that Deplete the Ozone Layer states (Article 6): “. . . the Parties

shall assess the control measures . . . on the basis of available scientific, environmental, technical, and economic information.” To provide the mechanisms whereby these assessments are conducted, the Protocol further states: “. . . the Parties

shall convene appropriate panels of experts” and “the panels will report their conclusions . . . to the Parties.”

To meet this request, the Scientific Assessment Panel, the Environmental Effects Panel, and the Technology and

Economic Assessment Panel have each prepared, about every 3-4 years, major assessment reports that updated the state

of understanding in their purviews. These reports have been scheduled to be available to the Parties in advance of their

meetings at which they will consider the need to amend or adjust the Protocol.

The Sequence of Scientific Assessments

The current 2002 report is the latest in a series of nine scientific assessments prepared by the world’s leading

experts in the atmospheric sciences and under the international auspices of the World Meteorological Organization

(WMO) and/or the United Nations Environment Programme (UNEP). This report is the fifth in the set of major assessments that have been prepared by the Scientific Assessment Panel directly as input to the Montreal Protocol process. The

chronology of all the scientific assessments on the understanding of ozone depletion and their relation to the international

policy process is summarized as follows:

Year Policy Process Scientific Assessment

1981 The Stratosphere 1981: Theory and Measurements. WMO No. 11.

1985 Vienna Convention Atmospheric Ozone 1985. Three volumes. WMO No. 16.

1987 Montreal Protocol

1988 International Ozone Trends Panel Report 1988.

Two volumes. WMO No. 18.

1989 Scientific Assessment of Stratospheric Ozone: 1989.

Two volumes. WMO No. 20.

1990 London Adjustments and

Amendment

1991 Scientific Assessment of Ozone Depletion: 1991. WMO No. 25.

1992 Methyl Bromide: Its Atmospheric Science, Technology, and

Economics (Montreal Protocol Assessment Supplement).

UNEP (1992).

1992 Copenhagen Adjustments

and Amendment

1994 Scientific Assessment of Ozone Depletion: 1994. WMO No. 37.

1995 Vienna Adjustment
x

PREFACE

Year Policy Process Scientific Assessment

1997 Montreal Adjustments

and Amendment

1998 Scientific Assessment of Ozone Depletion: 1998. WMO No. 44.

1999 Beijing Amendment

2002 Scientific Assessment of Ozone Depletion: 2002. WMO No. 47.

2003 15th Meeting of the Parties

The Current Information Needs of the Parties

The genesis of Scientific Assessment of Ozone Depletion: 2002 occurred at the 11th Meeting of the Parties to the

Montreal Protocol in Beijing, China, at which the scope of the scientific needs of the Parties was defined in their Decision

XI/17.5(a): “To request the Scientific Assessment Panel to include the following in the 2002 scientific assessment:

(a) An evaluation of the observed trends in controlled substances and their consistency with reported production

of ODS;

(b) A quantification of the ozone-depleting impacts of new (e.g., short-lived) halogen-containing substances;

(c) A characterization of methyl bromide sources and sinks and the likely quantitative implications of the results

for the ozone layer;

(d) A characterization of the known interrelations between ozone depletion and climate change including feedbacks between the two;

(e) A description and interpretation of the observed changes in global and polar ozone and in ultraviolet radiation, as well as set future projections and scenarios for those variables, taking into account also the expected

impacts of climate change...”.

The Assessment Process

The formal planning of the current assessment was started early in 2001. At the request of the Scientific Assessment

Panel, the Parties suggested experts from their countries who could participate in the process, and those suggestions contributed about half of the participants who served as authors, contributors, and reviewers. Furthermore, an ad hoc international scientific steering group also suggested participants from the world scientific community. In addition, this steering

group contributed to crafting the outline of the assessment report. As in previous assessments, the participants represented experts from the developed and developing world. The developing-world experts bring a special perspective to

the process, and their involvement in the process contributes to capacity building.

The information of the 2002 assessment is contained in five chapters, with most containing past trends and future

projections associated with ozone-layer topics: 

Chapter 1. Controlled Substances and Other Source Gases

Chapter 2. Very Short-Lived Halogen and Sulfur Substances

Chapter 3. Polar Stratospheric Ozone: Past and Future

Chapter 4. Global Ozone: Past and Future

Chapter 5. Surface Ultraviolet Radiation: Past and Future

The interactions between the ozone layer and the climate system are varied and appear appropriately as a special section

in most of the chapters.

A special resource for the Panel’s work was the earlier report, Aviation and the Global Atmosphere. This 1999

assessment of the impacts of aviation on ozone depletion and climate change was a collaboration of the Intergovernmental

Panel on Climate Change (IPCC) and the Scientific Assessment Panel of the Montreal Protocol. The assessment had

been requested by the International Civil Aviation Organization (ICAO). Because this comprehensive study had been
recently done, the present 2002 assessment could cite the major relevant findings of the 1999 study and provide any

updates of knowledge that had occurred.

The initial plans for the chapters of the 2002 Scientific Assessment Panel’s report were examined at a meeting that

occurred on 27-28 June 2001 in London, United Kingdom. The Lead Authors and Cochairs focused on the content of the

draft chapters and establishing the needs for coordination among the chapters.

The first drafts of the chapters were examined at a meeting that occurred on 28-30 November 2001 in Fairfax,

Virginia, United States, at which the Lead Authors, Cochairs, and a small group of international experts focused on the

scientific content of the draft chapters.

The second drafts of the chapters were reviewed by 133 scientists worldwide in a mail peer review. Those comments were considered by the authors. At a Panel Review Meeting in Les Diablerets, Switzerland, held on 24-28 June

2002, the responses to these mail review comments were proposed by the authors and discussed by the 74 participants.

Final changes to the chapters were decided upon at this meeting. The Executive Summary contained herein (and posted

on the UNEP and WMO web sites on 23 August 2002) was prepared and completed by the attendees of the Les Diablerets

meeting.

The 2002 State-of-Understanding Report

In addition to the scientific chapters and the Executive Summary, the assessment also focuses on a set of questions

that are frequently asked about the ozone layer. Based upon the scientific understanding represented by the assessments,

answers to these frequently asked questions were prepared, with different readerships in mind, e.g., students and the general public. These questions and answers are included in this report.

The final result of this two-year endeavor is the present assessment report. As the accompanying list indicates, the

Scientific Assessment of Ozone Depletion: 2002 is the product of 275 scientists from the developed and developing world

who contributed to its preparation and review1 (170 scientists prepared the report and 182 scientists participated in the

peer review process).

xi

PREFACE

1 Participating were Albania, Argentina, Armenia, Australia, Austria, Belgium,Bolivia, Brazil, Canada, Chile, Colombia, Denmark, Egypt, Estonia,

Finland, France, Germany, Greece, India, Iran, Italy, Japan, Kenya, Malaysia, New Zealand, Norway, Poland, Russia, South Africa, Sweden, Switzerland,

Taiwan R.O.C., The Netherlands, The People’s Republic of China, Togo, United Kingdom, United States of America, and Venezuela.
xiii

EXECUTIVE SUMMARY

The provisions of the 1987 Montreal Protocol on Substances that Deplete the Ozone Layer include the requirement

that the Parties to the Protocol base their future decisions on the current scientific, environmental, technical, and economic information that is assessed through panels drawn from the worldwide expert communities. To provide that input

to the decision-making process, advances in understanding on these topics were assessed in 1989, 1991, 1994, and 1998.

This information helped support discussions among the Parties that led to the subsequent Amendments and Adjustments

of the 1987 Protocol. The 2002 Scientific Assessment summarized here is the fifth in that series.

RECENT MAJOR FINDINGS AND CURRENT SCIENTIFIC UNDERSTANDING

Since the Scientific Assessment of Ozone Depletion: 1998, numerous laboratory investigations, atmospheric

observations, and theoretical and modeling studies have produced new key findings and have strengthened overall understanding of the ozone layer and its effect on ultraviolet (UV) radiation. These advances are highlighted in the following

summary of the current understanding of the impact of human activities and natural phenomena on the ozone layer and

the coupling of the ozone layer and the climate system.

Changes in Ozone-Depleting Compounds

• In the troposphere (i.e., lower atmosphere), observations show that the total combined effective abundance

of ozone-depleting compounds continues to decline slowly from the peak that occurred in 1992-1994. Total

chlorine is declining, while bromine from industrial halons is still increasing, albeit at a slower rate than was

occurring previously (and as reported in the 1998 Assessment). Total tropospheric chlorine from the long- and

short-lived chlorocarbons was about 5% lower in 2000 than that observed at its peak in 1992-1994, and the rate of

change in 2000 was about -22 parts per trillion (ppt) per year (-0.6% per year). The once-dominant influence of

methyl chloroform (CH3CCl3) on this total decline is diminishing because the atmospheric abundance of methyl

chloroform is sharply decreasing. Total chlorine from the major chlorofluorocarbons (CFCs) is no longer increasing,

in contrast to the slight increase that was occurring at the time of the 1998 Assessment. Specifically, in 2000, the

atmospheric abundances of CFC-11 and CFC-113 continued to decrease, while the rate of increase of CFC-12

had slowed. Total tropospheric bromine from halons continued to increase at about 3% per year, which is about

two-thirds of the rate for 1996 reported in the 1998 Assessment. The observed abundances of CFCs, hydrochlorofluorocarbons (HCFCs), and methyl chloroform in the lower atmosphere continue to be consistent with reported

production and estimated emissions.

• Analyses of air trapped in snow since the late 19th century have confirmed that non-industrial sources of the

CFCs, halons, and major chlorocarbons were insignificant. Since the previous Assessment, analyses of firn air

(i.e., air trapped in snow above glaciers) have revealed the abundance of long-lived atmospheric species at the time

the air became trapped. As a result, trends in the atmospheric abundance for many ozone-depleting substances have

been traced over the past century, to well before significant industrial sources of the compounds existed. These

records show that the mixing ratios of the CFCs, halons, carbon tetrachloride (CCl4), methyl chloroform, and

HCFCs in the oldest air sampled are negligible compared with the amounts measured in today’s background

atmosphere. Further, the deduced 20th century records for these compounds are broadly consistent with calculated

histories based on records of industrial production. The data suggest that substantial natural sources exist for

atmospheric methyl bromide (CH3Br). They also show increases throughout the 20th century, but these increases do

not allow unambiguous quantification of the industrial fraction of methyl bromide emissions in recent years. The

estimate of this fraction, based on an assessment of understanding of the budget of this gas, remains at 10-40%, as

given in the 1998 Assessment.
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xiv

• The abundances of HCFCs in the lower atmosphere continue to increase. HCFCs are among the gases used as

transition substitutes for CFCs, halons, and chlorinated solvents. In the year 2000, HCFCs represented 6% of total

chlorine abundance from anthropogenic gases in the lower atmosphere. The rate of increase in chlorine from

HCFCs was constant at 10 parts per trillion per year from 1996 to 2000.

• Observations in the stratosphere indicate that the total chlorine abundance is at or near a peak, while bromine

abundances are probably still increasing. The sum of hydrogen chloride (HCl) and chlorine nitrate (ClONO2) is

an effective surrogate for the abundance of stratospheric chlorine. An extended time series of ground-based measurements shows that the total stratospheric column amounts of these species, which have grown steadily for decades,

have plateaued in recent years. Further, space-based measurements of HCl in the upper stratosphere indicate a

broadly similar behavior. There are indications that bromine abundances in the stratosphere increased during the

1990s, but changes in stratospheric bromine are not as well characterized as those of stratospheric chlorine. These

stratospheric changes are consistent with expectations based on the understanding of trace-gas trends in the troposphere, stratospheric chemistry, and atmospheric transport from the troposphere to the stratosphere.

• Very short-lived organic chlorine-, bromine-, and iodine-containing source gases have the potential to deplete

stratospheric ozone, but quantitative estimation of their potentials is more challenging than for longer-lived

species like CFCs. The very short-lived compounds reside in the atmosphere for a few months or less because they

are rapidly decomposed chemically in the troposphere. Yet, a fraction of their emissions and the products from

their tropospheric destruction can potentially reach the stratosphere. For example, observations suggest that nonanthropogenic bromoform (CHBr3) produced largely in the oceans does make a non-negligible contribution to the

total stratospheric bromine abundance. The magnitude of the ozone depletion by very short-lived compounds will

depend critically on the location and season of their emissions and on the properties of their degradation products.

The traditional use of a single number for their Ozone Depletion Potential (ODP), which is possible for longer-lived

species, is therefore not directly applicable to the very short-lived species. Three-dimensional model simulations

also suggest that very short-lived compounds emitted in the tropics would be more readily transported to the stratosphere than those emitted at higher latitudes, thus leading to greater ozone loss for tropical emissions. ODP values

estimated by three-dimensional models are currently uncertain because of difficulties in modeling the complexities

of transport processes and the lack of data on the products of the tropospheric degradation. A recent study on npropyl bromide, one of the compounds proposed for possible future use, showed that for emissions that are uniform

over the global land masses away from the poles, roughly 0.5% of the bromine emitted as n-propyl bromide reaches

the stratosphere, resulting in an ODP of 0.04. Other ODP values reported in that study are up to 0.1 from tropical

emission, and values up to 0.03 and 0.02 for emissions restricted to north of 20°N and 30°N, respectively. Therefore,

the impact of very short-lived compounds can be significant if their emissions are large.

Changes in the Ozone Layer over the Poles and Globally

• Springtime Antarctic ozone depletion due to halogens has been large throughout the last decade. Since the

early 1990s, the minimum total column (i.e., overhead) ozone amount has been ~100 Dobson units (DU). The

monthly total column ozone amounts in September and October have continued to be about 40 to 50% below preozone-hole values, with up to a local 70% decrease for periods of a week or so. During the last decade, the average

ozone hole area in the spring has increased in size, but not as rapidly as during the 1980s. The area of the ozone

hole varies from one year to another, and it is not yet possible to say whether the area of the ozone hole has maximized. In recent years, the ozone hole has also persisted into early summer, increasing its impact on ultraviolet

radiation.

• In some recent cold Arctic winters during the last decade, maximum total column ozone losses due to halogens have reached 30%. Arctic winter/spring ozone loss is highly variable due to changes in stratospheric

meteorological conditions from one winter to another, but it is now better understood because of numerous new

observations and model comparisons. There is general agreement between analyses that quantify Arctic chemical

ozone loss for the 1999/2000 winter/spring season. That well-studied year was distinguished by persistent low temperatures, an ozone loss reaching 70% near 20 km, and total column ozone losses greater than 80 Dobson units

(~20-25%) by early spring. In contrast, during the warmer, more disturbed Arctic winter of 1998/1999, the
EXECUTIVE SUMMARY

xv

estimated chemical loss was very small. Three of the last four Arctic winters have been warm, with little ozone

loss; six of the previous nine winters were cold, with larger ozone losses.

• Ozone remains depleted in the midlatitudes of both hemispheres. The global-average total column ozone

amount for the period 1997-2001 was approximately 3% below the pre-1980 average values. Observed changes

occur primarily in midlatitudes and in polar regions; no significant trends in total column ozone have been observed

in the tropics (25°N-25°S). There are differences in ozone behavior between the two hemispheres. In particular,

the average amounts of total column ozone over the period 1997-2001 were 3% and 6% below the pre-1980 values

in the Northern Hemisphere midlatitudes (35°N-60°N) and the Southern Hemisphere midlatitudes (35°S-60°S),

respectively. The seasonality of total column ozone changes (1997-2001 relative to pre-1980) is different in the

Northern Hemisphere and Southern Hemisphere. Over Northern Hemisphere midlatitudes, the largest ozone

decreases are observed during winter/spring (~4%), with summer/autumn decreases approximately half as large.

Over Southern Hemisphere midlatitudes, long-term ozone decreases exhibit a similar magnitude (~6%) during all

seasons. 

• Models including observed changes in halocarbons, source gases, and aerosols (i.e., airborne fine particles)

capture the observed long-term ozone changes in northern and southern midlatitudes. The two-dimensional

assessment models also reproduce much of the interannual ozone variations in the midlatitudes of the Northern

Hemisphere, but do less well in the Southern Hemisphere. For example, observations show different ozone behavior

in the Northern and Southern Hemispheres following the major eruption of the Mt. Pinatubo volcano in the early

1990s, whereas models that include aerosol-enhanced, halocarbon-ozone chemistry suggest hemispherically symmetric ozone loss during the post-eruption period. Changes in dynamical processes help to explain some of the

ozone variations in Northern Hemisphere midlatitudes. They have also contributed to winter/spring trends in the

Northern Hemisphere. However, because chemical and dynamical processes are coupled, their contributions to

ozone changes cannot be assessed in isolation.

• Chemistry-climate models predict that springtime Antarctic ozone levels will be increasing by 2010 because

of projected decreases of halogens in the stratosphere. A return to pre-1980 total column ozone amounts in the

Antarctic is expected by the middle of this century.

• Arctic ozone depletion is highly variable and difficult to predict, but a future Arctic polar ozone hole similar

to that of the Antarctic appears unlikely. Low ozone, as seen in some recent years, can however be expected

again, and the Arctic stratosphere will be most vulnerable to other perturbations (for example, if there were to be an

increase in the abundance of stratospheric aerosols from volcanic eruptions) during the next decade or so. Sustained

very low Arctic ozone column amounts similar to those seen in the Antarctic are not predicted by the current

chemistry-climate models. Such extreme ozone depletion during the next decade or so, when halogen abundances

should still be close to their maximum, would require conditions that are unprecedented in about 40 years of

Northern Hemisphere meteorological observations and, therefore, are considered highly unlikely to occur in the

future. 

• The global ozone layer recovery is expected to be linked mainly to decreasing chlorine and bromine loading,

but other factors are likely to contribute. The expected decrease in the amount of stratospheric chlorine and

bromine over the next 50 years is predicted to lead to an increase in the global amount of total column ozone,

although there are differences in the projected rate of this increase predicted by different models. Stratospheric

cooling (due mainly to projected carbon dioxide (CO2) increases) is predicted to enhance the future ozone increase

in the upper stratosphere. However, a reliable assessment of this effect on total column ozone is limited by uncertainties in the lower stratospheric response to these changes. Changes in atmospheric transport are difficult to predict, and their impact on stratospheric ozone could be either positive or negative. Projected increases in methane

(CH4) and nitrous oxide (N2O) are predicted to have small chemical effects on the rate of increase of global total

column ozone in the next 50 years, but could become more significant later in the 21st century. Future changes in

the ozone in the lower atmosphere are highly dependent upon the scenario adopted for future emissions of ozone

precursors, but all scenarios adopted by the 2001 report of the Intergovernmental Panel on Climate Change (IPCC)

lead to predicted increases in tropospheric ozone up to 2050.
EXECUTIVE SUMMARY

xvi

Changes in Ultraviolet Radiation

• Changes in the duration and spatial extent of the ozone hole are more important for Antarctic surface ultraviolet (UV) radiation levels than the annual ozone minimum. Enhanced values of UV radiation continue to be

observed at high latitudes in the Southern Hemisphere under the Antarctic ozone hole. The highest biologically

weighted UV doses under the ozone hole are typically not observed in October when maximum ozone depletion

occurs, but in November and early December when solar elevations are higher and low ozone values are still

prevailing.

• Additional measurements continue to confirm that decreases in ozone column amounts lead to increases in

UV radiation. Calculations of UV irradiance based on relationships with total ozone and total irradiance (from

pyranometers) suggest that UV irradiance has increased since the early 1980s by 6-14% at more than 10 sites

distributed over mid- and high latitudes of both hemispheres. These results are consistent with spectral ultraviolet

irradiance measurements and with estimates from satellite measurements. The complicated spatial and temporal

distributions of the predominant variables that affect ultraviolet radiation at the surface (for example, clouds, airborne fine particles, snow cover, sea ice cover, and total ozone) continue to limit the ability to describe fully surface

ultraviolet radiation on the global scale, whether through measurements or model-based approaches. As was noted

in the previous Assessment, the spectral surface ultraviolet data records, which started in the early 1990s, are still

too short and too variable to permit the calculation of statistically significant long-term (i.e., multidecadal) trends.

The Ozone Layer and Climate Change

• The understanding of the impact of ozone depletion on climate change has been strengthened. There has been

a global and annual-mean cooling of the stratosphere over the past two decades, which can be largely attributed to

the observed stratospheric ozone depletion and increases in well-mixed greenhouse gases and water vapor. As has

been noted in past assessments, cooling of the lower stratosphere leads to cooling of the Earth’s climate system.

The vertical profile of ozone depletion in the lowermost stratosphere, which is an important factor in the magnitude

of the radiative forcing, is now more accurately estimated from additional years of observations with reduced volcanic perturbations. Averaged ozone depletion has remained close to that of the late 1990s over much of the world,

and therefore the recommended globally averaged radiative forcing of the climate system implied by this Assessment

is the same as that recommended by the 2001 IPCC Assessment. The stratospheric radiative forcing due to ozone

decreases since 1980 offsets about 20% of the positive forcing due to the increases in abundances of well-mixed

greenhouse gases over that same time period.

• Other atmospheric changes influence both the ozone layer and the climate system. Observations have provided stronger evidence for a widespread increase in stratospheric water vapor, which plays a role both in cooling

the lower stratosphere and in depleting ozone through chemical interactions, thereby contributing to climate

processes. However, the water vapor trends are not fully defined, nor are their cause understood. Methane, nitrous

oxide, and carbon dioxide are all important greenhouse gases, and all exert some influence on ozone depletion.

Further, surface ultraviolet radiation may be directly affected, both positively and negatively, by the effects of climate change (for example, changing cloudiness), making prediction of long-term changes in surface radiation

arising from all causes quite uncertain.

• New research has begun to explore the coupling between climate change and the recovery of the ozone layer.

A number of models have been run to explore the feedback between climate and the ozone layer. As noted earlier,

they have shown that past changes in ozone have contributed, together with well-mixed greenhouse gases, to a

cooling of the stratosphere. Future changes in well-mixed greenhouse gases will affect the future evolution of

ozone through chemical, radiative, and dynamic processes. In this highly coupled system, attribution is difficult;

studies are ongoing. Stratospheric cooling (due mainly to projected carbon dioxide increases) is predicted to

enhance future ozone amounts in the upper stratosphere. However, a reliable assessment of these effects on total

column ozone is limited by uncertainties in lower stratospheric response to these changes.
EXECUTIVE SUMMARY

xvii

ADDITIONAL SCIENTIFIC EVIDENCE AND RELATED INFORMATION

Halocarbon Abundances

• Trends of ozone-depleting substances in the atmosphere have been updated, and 20th century trends have been

deduced from firn air. In 2000, tropospheric mixing ratios of CFC-11 and CFC-113 declined faster than in 1996,

and mixing ratios of CFC-12 were still increasing, but more slowly. The rapid drop in global methyl chloroform

emission has led to an exponential decay in its mixing ratio since 1998; mixing ratios of this gas in 2000 were less

than one-half of the peak observed in 1992. The rate of decline observed for methyl chloroform during 2000 was

about two-thirds of what it was in 1996.

• The total effect of all ozone-depleting halogens in the atmosphere, as estimated by calculating chlorine equivalents

from atmospheric measurements of chlorine- and bromine-containing gases, continues to decrease. As of mid2000, equivalent organic chlorine in the troposphere was nearly 5% below the peak value in 1992-1994. The recent

decrease is slightly slower than in the mid-1990s, owing to the reduced influence of methyl chloroform on this

decline.

• Substantial reductions in the emissions of ozone-depleting substances during the 1990s as inferred from atmospheric measurements are consistent with controls on production and consumption in the fully amended and adjusted

Montreal Protocol. Consumption in developing countries is now a significant contributor to global emissions. The

year 1999 was the first in which production and consumption of a class of ozone-depleting substances (the CFCs)

was restricted in all Parties to the Montreal Protocol. Atmospheric measurements are consistent with emissions

derived from reported production data for CFCs.

• The updated, best-estimate scenario for future halocarbon mixing ratios suggests that the atmospheric burden of

halogens will return to the 1980 pre-Antarctic-ozone-hole levels around the middle of the 21st century, provided

continued adherence to the fully amended and adjusted Montreal Protocol. Only small improvements would arise

from further reduced production allowances in the future.

• Discrepancies reported in past assessments between atmospheric observations and expectations based on industryreported production and emissions have narrowed substantially for HCFC-142b. This improvement stems from a

better description of the functions relating emissions to usage in foam applications.

Halocarbon Lifetimes

• The global lifetime of carbon tetrachloride is estimated to be about 26 years, or about 25% shorter than in the previous (1998) Assessment. This shorter lifetime stems from identification of an ocean sink that is inferred from

widespread observations of carbon tetrachloride undersaturation in surface waters of the ocean. Emissions inferred

from atmospheric measurements and this lifetime are about 7 times greater than the limits to global production set

for 2005. 

• The lifetime of methyl chloroform has been revised from 4.8 years to 5.0 years based upon new observations. The

implications of this change on estimates of atmospheric hydroxyl (OH) suggest lifetimes up to 5% longer for

HCFCs, hydrofluorocarbons (HFCs), methane, and all other gases removed from the atmosphere by this important

oxidant. These changes affect the Global Warming Potentials (GWPs) and Ozone Depletion Potentials (ODPs) calculated for these gases.

Methyl Bromide, Methyl Chloride, and Halons

• Atmospheric histories inferred from Southern Hemisphere air archives and Antarctic firn air suggest that, assuming

similar changes have occurred in both hemispheres, the sum of organic bromine from methyl bromide (CH3Br) and

halons has more than doubled since the mid-1900s.
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xviii

• A substantial imbalance remains in estimates of source and sink magnitudes for both methyl bromide and methyl

chloride (CH3Cl); known sinks outweigh sources for both of these gases. New sources of methyl bromide from

individual crops and ecosystems have been identified, and new sources of methyl chloride from tropical plants have

been discovered. These findings have narrowed the budget imbalances for both of these gases.

• The best estimate for the global lifetime of methyl bromide remains at 0.7 (0.5-0.9) years. Additional studies

directly related to estimating loss processes for methyl bromide have narrowed the uncertainties slightly, but do not

suggest large revisions to this lifetime. The fraction of emissions derived from industrially produced methyl bromide is unchanged at 10-40% based upon the current understanding of source and sink magnitudes.

Very Short-Lived Ozone-Depleting Compounds

• Very short-lived natural and anthropogenic bromine and iodine source gases with surface concentrations of a few

parts per trillion (ppt) could make a non-negligible contribution to the current inorganic bromine and iodine budgets,

since the stratospheric concentrations of inorganic bromine and iodine are about 20 ppt and less than 1 ppt, respectively. The transport of inorganic bromine associated with very short-lived bromine source gases from the troposphere to the stratosphere may contribute to the stratospheric inorganic bromine budget.

• The most efficient route for transport of very short-lived substances and their degradation products from the surface

to the stratosphere is in the tropics. In the tropics, the vertical transport times from the boundary layer to the upper

troposphere are short, and air that enters the stratosphere through the tropical tropopause may remain in the stratosphere for a year or longer. A significant fraction of the emitted very short-lived substances can be expected to

reach the tropical tropopause layer because current estimates indicate that air at the base of the layer is replaced by

convection from the tropical boundary layer on a time scale of 10 to 30 days. A few percent of the air in the tropical

tropopause layer is expected to enter the stratosphere through the tropical tropopause. Other transport pathways

exist in the extratropics for the transfer of very short-lived substances and their degradation products to the extratropical lower stratosphere.

• The main uncertainties in estimating the impact of very short-lived source gases lie in the physical and dynamical

processes transporting these substances into the stratosphere and in the chemistry of their degradation products.

Given the complexity, three-dimensional numerical models are the preferred tools to evaluate the Ozone Depletion

Potential for very short-lived source gases. Significant uncertainties exist in the treatment of dynamical and physical processes in such models.

• Two model studies simulated the atmospheric distribution of bromoform (CHBr3), assuming a simplified ocean

source that is uniform over space and time. The results indicate that the ocean source causes an average surface

mixing of 1.5 ppt for bromoform and maintains about 1 ppt of bromine in the stratosphere. The simulation shows

that one-half to three-fourths of the bromine from bromoform enters the stratosphere in the form of inorganic degradation products.

• The Ozone Depletion Potential was calculated from three separate model studies for n-propyl bromide (n-PB,

CH3CH2CH2Br). Reaction with hydroxyl (OH) removes n-PB, with local photochemical lifetimes in the tropical

troposphere of about 10-20 days. Laboratory data, particularly on bromoacetone, indicate that n-propyl bromide

degradation products have lifetimes shorter than a couple of days. Two of the three modeling studies provided

values only for direct transport of n-PB to the stratosphere. The third study computed contributions from direct

transport and the transport of degradation products to the stratosphere. In the latter study, values of the Ozone

Depletion Potential are up to 0.1 for tropical emissions and 0.03 for emissions restricted to northern midlatitudes.

In both cases, about two-thirds of the effect is from the transport of degradation products to the stratosphere.

• Laboratory data on iodine chemistry have led to downward revision of the efficiency of iodine for depleting ozone

in the stratosphere. The revised estimated efficiency factor (~150-300) is still higher than that of bromine (~45).
EXECUTIVE SUMMARY

xix

Polar Ozone

ANTARCTIC

• Springtime Antarctic ozone depletion remains very large (with daily local total column values reaching 60-70%

less than pre-ozone-hole conditions), with minimum values of about 100 DU (Dobson units) seen every year since

the early 1990s. These observations reflect the almost complete ozone loss in the 12-20 km range and do not imply

that ozone recovery has begun. Such low ozone is consistent with current understanding of stratospheric chemistry

and dynamics.

• The area enclosed by the 220 DU contour (a measure of the severity of the ozone hole) shows an increase in recent

years, so that it is not yet possible to say that the ozone hole has reached its maximum. Much of the change appears

to be associated with processes at the edge of the polar vortex and is consistent with meteorological variability and

the almost-constant halogen loading.

• Observations show that the Antarctic polar vortex and the associated ozone hole persist later than during the 1980s.

Over the last decade, the vortex has generally broken up in late November to early December, in contrast to breakup

in early November during the 1980s.

• Satellite and radiosonde observations show that the springtime Antarctic lower stratosphere has cooled. During the

1979-2000 period, the linear cooling trend exceeded 1.5 K/decade at 70°S. Modeling studies reaffirm that ozone

loss is the major cause of the springtime cooling and the increased persistence of the Antarctic polar vortex. Wellmixed greenhouse gas increases contribute to the annually averaged cooling. Stratospheric water vapor increases

may also be contributing.

• Coupled chemistry-climate model simulations, which include the combined effects of the changes in halogens and

well-mixed greenhouse gases, broadly reproduce past trends in the total column ozone over the Antarctic. These

models suggest that the minimum column ozone occurs prior to 2010 and that recovery to 1980 levels may be

expected in the middle of the 21st century. The model response for the past and future changes is driven mainly by

the changes in stratospheric halogen loading, with ozone recovery occurring after the peak of halogen loading.

ARCTIC

• The magnitude of halogen-induced loss of ozone for all Arctic winters during the last decade has now been studied

with a variety of observationally based approaches. There is generally good agreement between different analyses

that quantify the chemical loss. In the 1999/2000 winter, for which the most comprehensive studies were conducted, agreement was better than 20% in the Arctic stratosphere at about 20 km.

• The Arctic winter/spring total column ozone amounts continue to show a high interannual variability, reflecting the

variable meteorology of the Northern Hemisphere stratosphere. Low column ozone amounts were present during

the cold winter of 1999/2000. That year was distinguished by persistent low temperatures, a local loss reaching

70% at 20 km, and column losses greater than 80 DU (~20-25%). In the warmer, more-disturbed winters of

1998/1999 and 2000/2001, very small ozone loss was observed. Three of the last four Arctic winters have been

warm, with little ozone loss; six of the previous nine winters were cold, with larger ozone losses.

• Significant chemical loss of ozone (~0.5 parts per million) in the lower stratosphere during January has been

observed in several cold Arctic winters, contributing about 25% to the overall loss of ozone over the winter. The

observations indicate that the loss occurred exclusively during periods when the air masses were exposed to sunlight. Nevertheless, these January ozone losses cannot be fully explained with the current understanding of the

photochemistry.

• Coupled chemistry-climate models capture the typical interannual variability of Arctic ozone levels. Because

Arctic temperatures are often near the threshold for polar stratospheric cloud (PSC) formation and hence the
EXECUTIVE SUMMARY

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initiation of perturbed chemistry, there is a strong sensitivity to model temperature biases of only a few degrees

Celsius. This places severe limits on the ability of the models to simulate past and predict future Arctic ozone

behavior in winter.

• A number of coupled chemistry-climate models run for this Assessment suggest that minimum Arctic ozone would

occur within the next two decades, the timing of which will depend on the meteorology. Low ozone, as seen in

some recent years, can be expected again, and the Arctic stratosphere will be most vulnerable to other perturbations

(e.g., aerosols from volcanic eruptions) during the next decade or so. Total column ozone amounts in the Arctic

similar to the extreme lows seen in the Antarctic are not predicted by these models (in contrast to earlier simpler

calculations considered in the 1998 Assessment). These extremely low values would require conditions that are

unprecedented in about 40 years of Northern Hemisphere meteorological observations.

• Satellite and radiosonde observations show that the springtime Arctic lower stratosphere has cooled. However,

becuae of large variability in the Arctic spring, the magnitude of the trend is uncertain there. A linear cooling trend

(exceeding 1.5 K/decade) is observed during the 1979-2000 period at 70°N. Modeling studies now suggest that

stratospheric ozone depletion has exerted an important influence on the springtime cooling of the Arctic lower stratosphere over the 1979-2000 period, but the degree of attribution is hindered by the large dynamical variability in

that region.

• Observations of bromine monoxide (BrO) in the winter Arctic vortex by in situ and remote detection techniques are

in broad agreement and are consistent with a total bromine budget of ~20 ± 4 parts per trillion. Modeling studies of

the latitudinal, seasonal, and diurnal variations in BrO column abundances agree well with observations from a

number of ground sites, indicating that the processes that govern bromine partitioning and its budget in the polar

regions are reasonably well understood.

• Bromine measurements now allow for more accurate assessment of the contribution of bromine to polar ozone loss.

At present, the fractional contribution of bromine to total ozone loss ranges between 30 and 60%, depending on

temperature and abundances of chlorine monoxide (ClO). Considering the observed leveling off of the strength of

sources of chlorine, the role of bromine in polar ozone loss will continue to increase relative to that of chlorine until

the current upward trends of the bromine source gases reverse.

• Removal of nitrogen compounds (denitrification) has been observed to occur in the Arctic lower stratosphere in

several cold winters. Removal of up to 70% of the total reactive nitrogen was observed at some levels of the lower

stratosphere in the winter of 1999/2000. Observations and modeling results show that denitrification in the 1999-

2000 Arctic lower stratosphere increased ozone loss by as much as 30% at 20 km in spring.

• The understanding of what causes denitrification has been improved considerably by the discovery in 1999-2000 of

large nitric-acid-containing particles (with diameters of 10 to 20 micrometers) in the Arctic polar lower stratosphere. Sedimentation of these particles can account for observed Arctic denitrification, although the mechanism

of formation of these sedimenting particles is uncertain. Therefore, sedimentation of ice containing dissolved nitric

acid, which has been the generally assumed mechanism in global stratospheric models, is not the dominant mechanism in the Arctic.

• The chemical composition of liquid and solid polar stratospheric cloud particles has been measured directly for the

first time. Most of the measured compositions are in agreement with model calculations for liquid particles and

nitric acid trihydrate, which have been used in stratospheric models for many years. These measurements improve

confidence in the particle types used in microphysical models that are central to simulations of polar ozone loss.
EXECUTIVE SUMMARY

xxi

Global Ozone

TOTAL COLUMN OZONE

• Global mean total column ozone for the period 1997-2001 was approximately 3% below the 1964-1980 average.

Since systematic global observations began, the lowest annually averaged global total column ozone occurred in

1992-1993 (about 5% below the pre-1980 average). These changes are evident in each available global dataset.

• No significant trends in total column ozone have been observed in the tropics (25°N-25°S) for 1980-2000. A

decadal variation of total column ozone (with peak-to-trough variations of ~3%) is observed in this region, approximately in phase with the 11-year solar cycle. Total column ozone trends become statistically significant in the

latitude bands 25°-35° in each hemisphere.

• There are a number of differences in total column ozone behavior between the two hemispheres:

- Averaged over the period 1997-2001, total column ozone in the Northern Hemisphere and Southern

Hemisphere midlatitudes (35°-60°) were about 3% and 6%, respectively, below their pre-1980 average values.

- The seasonality of total column ozone changes (1997-2001 relative to pre-1980) is different in the Northern

Hemisphere and Southern Hemisphere extratropics. Over Northern Hemisphere midlatitudes, larger ozone

decreases are observed during winter/spring (~4%), with summer/autumn decreases approximately half as

large. Over Southern Hemisphere midlatitudes, long-term ozone decreases exhibit a similar magnitude (~6%)

during all seasons.

- Pronounced negative anomalies are observed in the Northern Hemisphere midlatitudes time series during

1992-1995 in the winter/spring seasons. Similar anomalies are not seen in the Southern Hemisphere midlatitudes.

- There is a sharp drop in ozone at Southern Hemisphere midlatitudes during 1985-1986. A similar drop is not

observed in the Northern Hemisphere.

VERTICAL OZONE DISTRIBUTION

• Ozone profile trends derived from the Stratospheric Aerosol and Gas Experiment (SAGE) satellite instrument show

significant negative trends over latitudes 60°N to 60°S for altitudes ~35-50 km (with extremes near 40 km). Trend

maxima of –7 to –8%/decade over the period 1979-2000 are observed in the 35°-60° latitude bands of both hemispheres, with no significant interhemispheric differences. These satellite results are in good agreement with independent Umkehr ozone measurements over the Northern Hemisphere midlatitudes.

• The updated SAGE data reveal significant negative trends extending throughout the tropics in the small amount of

ozone above 30 km, a feature not observed in previous assessments based on shorter time records.

• The observed ozone depletion in the upper stratosphere is consistent with observed changes in anthropogenic chlorine. The vertical and latitudinal profiles of trends in the upper stratosphere are reproduced by photochemical

models, but the magnitude of changes are sensitive to concurrent trends in temperature and methane (CH4).

• Long-term ozonesonde measurements are primarily available for the Northern Hemisphere midlatitudes. Whereas

ozone between 20 and 27 km decreased continuously during 1980-2000, ozone between 10 and 20 km decreased

through the early 1990s and was relatively constant thereafter. This behavior is consistent with observed changes in

Northern Hemisphere midlatitude column ozone.
EXECUTIVE SUMMARY

xxii

OZONE-RELATED CONSTITUENTS

• Stratospheric aerosol variability over the past 25 years has been dominated by the effects of episodic volcanic eruptions, with subsequent recovery. Following the large eruption of Mt. Pinatubo in 1991, relaxation to a nonvolcanic

level continued to at least 1999. There is currently no evidence of a trend in the nonvolcanic aerosol loading.

• Stratospheric water vapor measurements at a single location (Boulder, Colorado, U.S., 40°N) for the period 1981-

2000 show a statistically significant increase of approximately 1%/year over altitudes 15-28 km. For the shorter

period 1991-2001, global satellite measurements covering latitudes 60°N-60°S show a similar trend of 0.6-0.8%/year

for altitudes ~25-50 km, but no significant trend at lower altitudes. The increases in water vapor are substantially

larger than can be explained by tropospheric methane trends. Characterization of stratospheric water vapor trends

is limited by the lack of global long-term measurements.

• Stratospheric column nitrogen dioxide (NO2) measurements from Lauder, New Zealand (45°S), for 1981-2000 and

Jungfraujoch, Switzerland (46°N), for 1985-2001 show statistically significant positive trends of approximately

5%/decade. There are also transient decreases observed after the El Chichón and Mt. Pinatubo eruptions, which are

broadly simulated by models that include heterogeneous chemistry on sulfate aerosols.

STRATOSPHERIC TEMPERATURE

• Observations indicate that, on an annual- and global-mean basis, the stratosphere has cooled over the last two

decades. In the lower stratosphere, global and annual mean temperatures for the late 1990s are approximately 1 K

lower than values in the late 1970s. Significant annual-mean cooling of the lower stratosphere over the past two

decades is found over midlatitudes of both hemispheres (approximately 0.6 K/decade), but no significant trends are

observed near the equator. The annual-mean temperature trends in the upper stratosphere are larger, with an approximately globally uniform cooling over 1979-1998 of about 2 K/decade near the stratopause (~50 km).

• Modeling studies indicate that changes in ozone, well-mixed greenhouse gases, and stratospheric water vapor can

explain the major features of the observed global and annual-mean stratospheric cooling over the past two decades.

Cooling due to ozone depletion dominates over the impact of well-mixed greenhouse gases in the lower stratosphere, while upper stratospheric temperature trends are due, roughly equally, to ozone and well-mixed greenhouse

gas changes.

ATTRIBUTION OF PAST CHANGES IN OZONE

• The vertical, latitudinal, and seasonal characteristics of changes in midlatitude ozone are broadly consistent with

the understanding that halogens are the primary cause, in line with similar conclusions from the 1998 Assessment.

• Assessment models forced by observed changes in halocarbons, source gases, and aerosols broadly reproduce the

long-term changes observed in midlatitude total column ozone (35°N-60°N and 35°S-60°S) from 1980 to 2000,

within the uncertainties of the observations and model range. However, the range of model results is large over

Southern Hemisphere midlatitudes, which is at least partly due to their differing treatments of the Antarctic ozone

hole. In addition, models suggest that the chemical signal of ozone loss following the major eruption of the Mt.

Pinatubo volcano in the early 1990s should have been symmetric between the hemispheres, but observations show

a large degree of interhemispheric asymmetry in midlatitudes.

• There is increased evidence that observed changes in atmospheric dynamics have had a significant influence on

Northern Hemisphere midlatitude column ozone on decadal time scales. Natural variability, changes in greenhouse

gases, and changes in column ozone itself are all likely to contribute to these dynamical changes. Furthermore,

because chemical and dynamical processes are coupled, their contributions to ozone changes cannot be assessed in

isolation.
EXECUTIVE SUMMARY

xxiii

FUTURE OZONE CHANGES

• The expected decrease in stratospheric chlorine loading over the next 50 years is predicted to lead to an increase in

the global total column ozone, although there are differences in the rate of increase between different twodimensional assessment models. Future ozone levels will also be influenced by other changes in atmospheric

composition and by climate change. Because of year-to-year variability, it could take as long as a decade to demonstrate a leveling of total column ozone.

• Stratospheric cooling (due mainly to projected CO2 increases) and the chemical influence of stratospheric methane

increases are predicted to enhance future ozone increases in the upper stratosphere. However, a reliable assessment

of these effects on total column ozone is limited by uncertainties in lower stratospheric response to these changes.

• Projected increases in methane (CH4) and nitrous oxide (N2O) (from the scenarios of the Intergovernmental Panel

on Climate Change, 2001) are predicted to have small effects on the rate of increase of global column ozone in the

next 50 years, when chlorine changes are the dominant effect. After that time, changes in CH4 and N2O become

relatively more important.

Ultraviolet Radiation

• Annually averaged erythemal irradiance, as reconstructed from pyranometer (total irradiance), total ozone, and

other meteorological measurements, increased by about 6-14% over the last 20 years at several mid- to highlatitude sites. Pyranometer and other meteorological data serve as proxies for parameters, other than ozone, that

affect ultraviolet (UV) radiation. At some sites approximately half of the changes can be attributed to total ozone

changes. These reconstructions are not UV measurements, and they contain several assumptions on the nature of

radiative transfer. The reconstructions should be not be considered to be representative on a global scale. It is

believed that the increases of UV irradiance derived from the ground-based reconstructed data are clear indicators

of the long-term changes that have occurred since the 1980s.

• There is clear evidence that the long-term UV changes are not driven by ozone alone, but also by changes in cloudiness, aerosols, and surface albedo. The relative importance of these factors depends on the local conditions. Results

from studies using ground-based and aircraft instruments suggest that the influence of tropospheric aerosols on UV

irradiance may be larger than previously thought and may affect large areas of the globe.

• UV increases associated with the ozone decline have been observed by spectral measurements at a number of sites

in Europe, North and South America, Antarctica, and New Zealand. Episodes of elevated UV irradiance associated

with low total column ozone amounts continue to occur in spring in mid-to-high latitudes.

• Satellite estimates of surface UV radiation from the Total Ozone Mapping Spectrometer (TOMS) dataset have been

compared with ground-based measurements at several more sites since the previous Assessment. In general the

estimates capture short-term and long-term variability. However, the estimates are systematically higher than

ground-based measurements at many sites. The differences in monthly average erythemal UV irradiance range

from about 0% at some clean sites to 40% at one site in the Northern Hemisphere. The fact that the agreement is

better at the cleaner sites suggests that the differences are caused by aerosols and/or pollutants near the ground.

New UV maps that include additional influencing parameters (e.g., cloud cover and albedo) derived from other

satellite data, when taken together with TOMS or Global Ozone Monitoring Experiment (GOME) ozone data, yield

better agreement with ground-based data.

• In the Antarctic, ozone depletion has been the dominant factor for increases in UV irradiance. The future evolution

of UV radiation is therefore expected to follow the ozone recovery. However, because of changes in other influencing factors, such as changes in cloud cover, aerosols, or snow/ice cover, UV radiation may not return exactly to

pre-ozone-hole values.
EXECUTIVE SUMMARY

xxiv

• Elsewhere, including the Arctic, the impact on UV radiation of other influencing factors can be comparable to the

impact of ozone depletion. The large uncertainties in future changes of these other factors prevent reliable predictions on the future evolution of UV irradiance. Furthermore, climate-change-induced trends in cloudiness and

snow/ice cover are expected to be seasonally and geographically dependent, leading to differences in future UV

irradiance in different parts of the world.

• A reanalysis of TOMS satellite data with respect to the influence of changes in cloudiness over Europe has confirmed that UV increases due to ozone depletion are partly masked by the increased cloudiness in some regions.

IMPLICATIONS FOR POLICY FORMULATION

The results from over three decades of research have provided a progressively better understanding of the interaction of humankind and the ozone layer. New policy-relevant insights into the roles of ozone-depleting gases have been

conveyed to decisionmakers through the international state-of-understanding assessment process. The research findings

in the Scientific Assessment of Ozone Depletion: 2002 that are summarized above are direct current scientific input to

governmental, industrial, and policy decisions associated with protection of the ozone layer.

• The Montreal Protocol is working, and the ozone-layer depletion from the Protocol’s controlled substances is

expected to begin to ameliorate within the next decade or so. The effectiveness of the Protocol is and will be

shown by several indicators. Global observations show that the total combined effective abundances of anthropogenic chlorine-containing and bromine-containing ozone-depleting gases in the lower atmosphere (troposphere)

peaked in the 1992-1994 time period and are continuing to decline. Furthermore, observations indicate that the

stratospheric abundances of ozone-depleting gases are now at or near a peak. Thereafter, stratospheric ozone

should increase, all other influences assumed constant, but ozone variability will make detection of the onset of the

long-term recovery difficult. For example, based on assumed compliance with the amended and adjusted Protocol

by all nations, the Antarctic ozone “hole,” which was first discerned in the early 1980s, is predicted to disappear by

the middle of this century—again with all other influences assumed constant.

• The ozone layer will remain particularly vulnerable during the next decade or so, even with full compliance.

With the atmospheric abundances of ozone-depleting substances being near their highest, the human-influenced

perturbations will be at or near their largest. Relative to the pre-ozone-hole abundances of 1980, the 1997-2001

losses in total column (i.e., overhead) ozone amounts are:

- about 4% at northern midlatitudes in winter/spring;

- about 2% at northern midlatitudes in summer/fall; and

- about 6% at southern midlatitudes on a year-round basis.

Calculations yield that such changes in ozone correspond to increases in surface erythemal radiation of at least 5, 2,

and 7%, respectively, if other influences such as clouds remain constant. In Antarctica, the monthly total column

ozone in September and October has continued to be about 40 to 55% below the pre-ozone-hole values, with up to a

local 70% decrease for periods of a week or so. Arctic ozone is highly variable. Estimates of the cumulative

winter/spring losses in the total column ozone amounts during the last 4 years range up to about 25%. Calculations

of corresponding increases in surface erythemal radiation are about 70 to 150% in the Antarctic springtime, with up

to 300% increases for the short-lived local ozone decreases. In the Arctic winter/spring, the corresponding calculated increases are up to 40%. Furthermore, if there were to be an increase in the abundance of stratospheric particles from a major volcanic eruption like that of Mt. Pinatubo in 1991, then the peak losses in total column ozone

and the increases in ultraviolet radiation could be larger. In the highly variable Arctic, larger depletion would be

expected if an unusually and persistently cold Arctic stratospheric winter like that of the 1999/2000 winter/spring

were to occur; conversely, smaller depletions are expected in particularly warm years.
EXECUTIVE SUMMARY

xxv

• Approaches to accelerating the recovery of the ozone layer are limited. This Assessment has made hypothetical

estimates of the upper limits of improvements that could be achieved if global anthropogenic production of ozonedepleting substances were to stop in 2003 or if global anthropogenic emissions of ozone-depleting substances were

to stop in 2003. Specifically: 

Production. Relative to the current control measures (Beijing, 1999) and recent production data, the equivalent

effective stratospheric chlorine loading above the 1980 level, integrated from 2002 until the 1980 level is reattained

(about 2050), could be decreased by the following amounts:

- 5%, if production of hydrochlorofluorocarbons (HCFCs) were to cease in 2003.

- 4%, if production of chlorofluorocarbons (CFCs) were to cease in 2003.

- 4%, if production of methyl bromide were to cease in 2003.

- 1%, if production of halons were to cease in 2003.

- 0.3%, if production of methyl chloroform were to cease in 2003.

These percentages would be about a factor of 2 smaller if the decreases were compared with the loading integrated

from 1980, which is when significant ozone depletion was first detected. A hypothetical elimination of all anthropogenic production of all ozone-depleting substances would advance the return of stratospheric loading to the pre1980 values by about 4 years.

Emissions. Similarly, the equivalent effective stratospheric chlorine loading above the 1980 level, integrated from

2002 until the 1980 level is reattained (about 2050), could be decreased by the following amounts:

- 11%, if emissions of halons were to cease in 2003.

- 9%, if emissions of chlorofluorocarbons (CFCs) were to cease in 2003.

- 9%, if emissions of hydrochlorofluorocarbons (HCFCs) were to cease in 2003.

- 4%, if emissions of methyl bromide were to cease in 2003.

- 3%, if emissions of carbon tetrachloride were to cease in 2003.

- 2%, if emissions of methyl chloroform were to cease in 2003.

Again, these percentages would be about a factor of 2 smaller if the decreases were compared with the loading integrated from 1980, which is when significant ozone depletion was first detected. A hypothetical elimination of all

emissions derived from industrial production of all ozone depleting substances would advance the return of stratospheric loading to the pre-1980 values by about 10 years.

• Failure to comply with the Montreal Protocol would delay or could even prevent recovery of the ozone layer.

For example, continued constant production of ozone-depleting substances at the 1999 amount would likely extend

the recovery of the ozone layer well past the year 2100. The total atmospheric abundance of ozone-depleting gases

will decline to pre-Antarctic-ozone-hole amounts only with adherence to the Montreal Protocol’s full provisions on

production of ozone-depleting substances.

• Estimating the impacts of very short-lived ozone-depleting substances on depletion of the ozone layer requires

new approaches, and, as requested by the Parties, this Assessment has described one such scientific approach.

The traditional concept of a single-valued Ozone Depletion Potential (ODP) is not directly applicable for these very

short-lived ozone-depleting substances, because their impacts on the ozone layer will depend on the season and

location of their emissions. These impacts would need to be assessed on a case-by-case basis, taking into account

how much, when, and where they are emitted. Such estimates can provide insight into the stratospheric contribution of natural emissions of these very short-lived substances (for example, bromoform) and can provide scientific

input into decisions associated with their industrial production/uses (for example, n-propyl bromide).

• The issues of ozone depletion and climate change are interconnected. The ozone-depletion phenomenon and

the greenhouse-warming phenomenon share many common chemical and physical processes. For example, as the

atmospheric abundances of the CFCs decline because of the Montreal Protocol’s provisions, their greenhousewarming contributions will decline. On the other hand, use of hydrofluorocarbons (HFCs) and HCFCs as substitutes for CFCs would cause the greenhouse-warming contributions of these new compounds to increase. Indeed,
EXECUTIVE SUMMARY

xxvi

global observations of many HFCs and HCFCs, as well as of hydrogen fluoride, confirm that these contributions

are currently increasing. As other examples, potential decisions associated with methane, nitrous oxide, and carbon

dioxide stemming from their greenhouse roles will also have direct and indirect effects on stratospheric ozone.

And, because ozone depletion acts to cool the climate system, recovery of the ozone layer over coming decades

would tend to warm the climate system.
CHAPTER 1

Controlled Substances and Other Source Gases

Lead Authors:

S.A. Montzka

P.J. Fraser

Coauthors:

J.H. Butler

P.S. Connell

D.M. Cunnold

J.S. Daniel

R.G. Derwent

S. Lal

A. McCulloch

D.E. Oram

C.E. Reeves

E. Sanhueza

L.P. Steele

G.J.M. Velders

R.F. Weiss

R.J. Zander

Contributors:

S.O. Andersen

J. Anderson

D.R. Blake

M.P. Chipperfield

E. Dlugokencky

J.W. Elkins

A. Engel

D. Harper

E. Mahieu

K. Pfeilsticker

J.-P. Pommereau

J.M. Russell III

G. Taylor

M. Van Roozendael

D.W. Waugh

CHAPTER 1

CONTROLLED SUBSTANCES AND OTHER SOURCE GASES

Contents

SCIENTIFIC SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.1

1.1 INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.5

1.2 HALOGENATED OZONE-DEPLETING GASES IN THE ATMOSPHERE . . . . . . . . . . . . . . . . . . . . . . . . . . 1.5

1.2.1 Updated Atmospheric Observations of Ozone-Depleting Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.5

1.2.1.1 Chlorofluorocarbons (CFCs) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.5

1.2.1.2 Halons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.9

1.2.1.3 Carbon Tetrachloride (CCl4) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.9

1.2.1.4 Methyl Chloroform (CH3CCl3) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.11

1.2.1.5 Hydrochlorofluorocarbons (HCFCs) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.11

1.2.1.6 Methyl Bromide (CH3Br) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.13

1.2.1.7 Methyl Chloride (CH3Cl) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.13

1.2.2 Twentieth Century Atmospheric Histories for Halocarbons from Firn Air . . . . . . . . . . . . . . . . . . . . 1.13

1.2.3 Total Atmospheric Chlorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.14

1.2.3.1 Total Organic Chlorine in the Troposphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.14

1.2.3.2 Total Inorganic Chlorine in the Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.15

1.2.3.3 Hydrogen Chloride (HCl) at 55 km Altitude in the Stratosphere . . . . . . . . . . . . . . . . . . . . 1.15

1.2.3.4 Stratospheric Chlorine Compared with Tropospheric Chlorine . . . . . . . . . . . . . . . . . . . . . 1.15

1.2.4 Total Atmospheric Bromine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.16

1.2.4.1 Total Organic Bromine in the Troposphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.16

1.2.4.2 Total Inorganic Bromine in the Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.17

1.2.4.3 Stratospheric Bromine Compared with Tropospheric Bromine . . . . . . . . . . . . . . . . . . . . . 1.18

1.2.5 Effective Equivalent Chlorine and Effective Equivalent Stratospheric Chlorine . . . . . . . . . . . . . . . 1.18

1.2.6 Fluorine in the Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.19

1.3 HALOCARBON SOURCES ESTIMATED FROM INDUSTRIAL PRODUCTION . . . . . . . . . . . . . . . . . . 1.19

1.3.1 CFCs and HCFCs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.21

1.3.2 Halons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .1.21

1.3.3 Carbon Tetrachloride (CCl4) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.21

1.3.4 Methyl Chloroform (CH3CCl3) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.21

1.4 HALOCARBON LIFETIMES, OZONE DEPLETION POTENTIALS, AND GLOBAL

WARMING POTENTIALS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.21

1.4.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.21

1.4.2 Lifetimes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.22

1.4.3 Fractional Release Factors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.27

1.4.4 Ozone Depletion Potential . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.28

1.4.5 Global Warming Potential . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.31

1.4.5.1 Radiative Forcing Updates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.31

1.4.5.2 Direct Global Warming Potentials . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.34

1.4.5.3 Net Global Warming Potentials . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.34

1.4.6 Hydroxyl Radical Changes with Time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.34

1.4.7 Other Sinks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.36

1.5 METHYL BROMIDE AND METHYL CHLORIDE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.36

1.5.1 Methyl Bromide (CH3Br) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.36

1.5.1.1 Anthropogenic and Anthropogenically Influenced Sources . . . . . . . . . . . . . . . . . . . . . . . . 1.37

1.5.1.2 Surface Interactions: Natural Sources and Sinks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.38
1.5.1.3 Atmospheric Lifetime . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.39

1.5.1.4 A Twentieth Century, Southern Hemispheric Trend and Its Implications . . . . . . . . . . . . . . 1.39

1.5.1.5 Atmospheric Budget . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.40

1.5.2 Methyl Chloride (CH3Cl) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.40

1.5.2.1 Atmospheric Distribution and Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.41

1.5.2.2 Sources of Atmospheric Methyl Chloride . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.41

1.5.2.3 Sinks of Atmospheric Methyl Chloride . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.43

1.5.2.4 Atmospheric Budget . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .1.43

1.6 ATMOSPHERIC HALOCARBON OBSERVATIONS COMPARED WITH EXPECTATIONS . . . . . . . . . . 1.44

1.6.1 Consistency Between Atmospheric Halocarbon Measurements and Known Sources and Sinks . . . 1.44

1.6.1.1 CFCs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.44

1.6.1.2 Methyl Chloroform . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.44

1.6.1.3 Carbon Tetrachloride . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.44

1.6.1.4 HCFCs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.47

1.6.1.5 Halons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.47

1.6.1.6 Regional Emissions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.47

1.6.2 Are Atmospheric Measurements Consistent with Compliance with the Fully 

Amended Montreal Protocol? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.47

1.7 OTHER TRACE GASES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.51

1.7.1 Carbon Dioxide (CO2) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.51

1.7.2 Methane (CH4) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.53

1.7.3 Nitrous Oxide (N2O) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.55

1.7.4 Carbon Monoxide (CO) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.56

1.7.4.1 Atmospheric Global Distribution and Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.56

1.7.4.2 CO Sources and Sinks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.56

1.7.4.3 Isotopic Data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.56

1.7.5 Carbonyl Sulfide (COS) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.57

1.7.6 Hydrofluorocarbons (HFCs) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.57

1.7.6.1 HFC-134a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.57

1.7.6.2 HFC-23 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.57

1.7.6.3 HFC-125, HFC-143a, and HFC-152a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.59

1.7.7 Fluorocarbons (FCs), SF6, and SF5CF3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.59

1.7.7.1 CF4 and C2F6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.59

1.7.7.2 C4F8 and C3F8 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.59

1.7.7.3 Sulfur Hexafluoride (SF6) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.61

1.7.7.4 Trifluoromethylsulfurpentafluoride (SF5CF3) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.61

1.8 HALOGENATED SOURCE GASES IN THE FUTURE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.61

1.8.1 Development of the Baseline Scenario Ab . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.61

1.8.2 Projecting Future Halocarbon Mixing Ratios . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.62

1.8.2.1 CFCs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.62

1.8.2.2 Methyl Chloroform . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.62

1.8.2.3 Carbon Tetrachloride . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.62

1.8.2.4 Halons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.67

1.8.2.5 HCFCs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.68

1.8.2.6 Methyl Bromide and Methyl Chloride . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .1.68

1.8.2.7 Effective Equivalent Chlorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.68

1.8.3 Scenario Results: Calculated Mixing Ratios and EESC . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.68

1.8.4 Uncertainty in the Baseline Scenario . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.70

1.8.5 Potential Influence of Future Climate Change on Halogenated Source Gases . . . . . . . . . . . . . . . . . 1.71

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.71
SCIENTIFIC SUMMARY

• As a result of the Montreal Protocol and its Amendments and Adjustments, organic chlorine in the global troposphere continues to decline slowly, and inorganic chlorine in the global stratosphere has stabilized.

(a) Total organic chlorine from long- and short-lived chlorocarbons continues to decline in the global troposphere.

Total tropospheric chlorine in 2000 was about 5% lower than observed at its peak (3.7 ± 0.1 parts per billion

(ppb)) in 1992-1994, and the rate of change in 2000 was about -22 parts per trillion per year (ppt yr-1

) (-0.6%

yr-1). The influence of methyl chloroform on this decline is diminishing. Total organic chlorine from chlorofluorocarbons (CFCs) is no longer increasing at Earth’s surface.

(b) Total inorganic chlorine in the atmosphere, as estimated from hydrogen chloride and chlorine nitrate total

column absorbance measurements, stopped increasing in 1997-1998 and has remained fairly constant since.

Because most atmospheric hydrogen chloride and chlorine nitrate reside in the stratosphere, this result provides

an estimate of chlorine changes in the stratosphere. These stratospheric changes are consistent with expectations based on our understanding of trace gas trends in the troposphere, stratospheric chemistry, and atmospheric mixing processes.

(c) Space-based measurements show that the global mean growth rate of hydrogen chloride at 55 km has been substantially less since 1997 than it was before that time. Although the measurements since 1997 also show shortterm variations, the recent data do not show the steady increases observed in the early 1990s.

• Recent results from observations of tropospheric and stratosphere bromine show the following:

(a) As of 2000, total organic bromine from halons continued to increase in the troposphere at about 0.2 ppt Br yr-1

(pmol mol-1 yr-1), and halons accounted for nearly 8 ppt of bromine. Global trend data for methyl bromide, a

gas that is responsible for about half of the 20 ppt of total bromine in today’s stratosphere, have not been updated

since the previous Assessment. Atmospheric histories inferred from Southern Hemisphere air archives and

Antarctic firn air suggest that, assuming similar changes in both hemispheres, total organic bromine from the

sum of methyl bromide and halons has more than doubled since the mid-1900s.

(b) Measurements of inorganic bromine in the stratosphere indicate a rate of increase consistent with observed tropospheric trends of halons and methyl bromide, but mixing ratios that are 4-6 ppt higher. Additional stratospheric bromine stems from the transport of nonanthropogenic, very short-lived gases (such as bromoform and

dibromomethane) and their degradation products to the stratosphere.

• Ozone-depleting halogens in the troposphere, as assessed by calculating chlorine equivalents from measurements

of organic chlorine- and bromine-containing source gases, continue to decrease. As of mid-2000, equivalent organic

chlorine in the troposphere was nearly 5% below the peak value in 1992-1994. The recent rate of decrease is

slightly less than in the mid-1990s owing to the reduced influence of methyl chloroform on this decline.

• Trends of ozone-depleting substances in the atmosphere have been updated, and 20th century trends have been

deduced from air trapped in snow above glaciers (firn air).

(a) In 2000, tropospheric mixing ratios of CFC-11 and -113 were decreasing faster than in 1996, and mixing ratios

of CFC-12 were still increasing, but more slowly.

(b) Global methyl chloroform mixing ratios have been declining exponentially since 1998 because of the rapid

drop in emissions to low levels; mixing ratios in 2000 were less than one-half of the peak observed in 1992. As

a result, the rate of decline observed for methyl chloroform (and chlorine from methyl chloroform) during 2000

was about two-thirds of what it was in 1996.

(c) Newly reported measurements of air from firn allow inferences regarding 20th century histories of ozonedepleting substances in the atmosphere. These data confirm that nonanthropogenic sources of chlorofluorocarbons, halons, carbon tetrachloride, methyl chloroform, and hydrochlorofluorocarbons are insignificant. They

are consistent, however, with there being substantial natural emissions of both methyl chloride and methyl

bromide.

SOURCE GASES

1.1
• Atmospheric halocarbon measurements can provide some assessment of past global compliance with production

restrictions in the Montreal Protocol, and these considerations provide the foundation for projecting halocarbon

mixing ratios in future scenarios.

(a) The substantial reductions in emissions of ozone-depleting substances during the 1990s that are inferred from

measured atmospheric trends are consistent with controls on production and consumption in the fully amended

and adjusted Montreal Protocol. Consumption in developing countries is now a significant contributor to

global emissions. The year 1999 is the first in which production and consumption of a class of ozone-depleting

substances (the CFCs) were restricted in all Parties to the Montreal Protocol. Atmospheric measurements are

consistent with emissions derived from reported global production data for CFCs.

(b) The updated, best-estimate scenario (Ab) for future halocarbon mixing ratios suggests that the atmospheric

burden of halogens will return to the 1980, pre-Antarctic-ozone-hole levels around the middle of this century,

provided continued adherence to the fully amended and adjusted Montreal Protocol. Only small improvements

would arise from reduced production allowances in the future. Lack of compliance to the Protocol controls

would delay or prevent recovery of stratospheric ozone.

• With respect to hydrochlorofluorocarbons (HCFCs) and hydrofluorocarbons (HFCs), the gases used as interim substitutes for CFCs, halons, and chlorinated solvents, this Assessment found the following:

(a) Organic chlorine from the HCFCs in the troposphere reached nearly 180 ppt in 2000 and represented 6% of

total chlorine from anthropogenic gases. The rate of increase in organic chlorine from HCFCs remained constant at about 10 ppt yr-1 from 1996 to 2000.

(b) Discrepancies reported in past Assessments between atmospheric observations and expectations based on

industry-reported production and emissions have narrowed substantially for HCFC-142b. This improvement

stems from a better description of the functions relating emissions to usage in foam applications.

(c) Mixing ratios of HFC-134a and -23 have continued to increase in recent years, and by 2000 each had approached

about 15 ppt in the background atmosphere. Three additional HFCs (HFC-125, -124, and -152a) have been

identified in the remote troposphere, but their abundance in 2000 was low (1-3 ppt).

• New lifetime recommendations are made in this Assessment:

(a) The global lifetime of carbon tetrachloride is estimated to be 26 years, or about 25% shorter than in the previous Assessment. This shorter lifetime stems from identification of an ocean sink that is inferred from widespread observations of carbon tetrachloride undersaturation in surface waters of the ocean. Emissions inferred

from this shorter lifetime and measured trends in 1996 are about a factor of 2 larger than those estimated from

industry production data for that year. This apparent discrepancy, however, is within the rather large uncertainties in both estimates. Emissions inferred from atmospheric measurements and a lifetime of 26 years are about

7 times greater than the limits to global production set for 2005.

(b) In this Assessment, the global lifetime of Halon-1211 is taken to be 16 years based upon the mid-range of modeling results. This lifetime estimate and estimates of emission magnitudes and atmospheric mixing ratios contain substantial uncertainties. Observational studies and emission histories have not reduced the uncertainties

in the global lifetime of 9-25 years calculated in models.

(c) The lifetime of methyl chloroform has been revised from 4.8 to 5.0 years based upon new observations. The

implications of this change on our estimates of atmospheric hydroxyl suggest slightly longer lifetimes for

HCFCs, HFCs, methane, and all other gases removed from the atmosphere by this important oxidant.

• With respect to methyl bromide and methyl chloride, ozone-depleting gases with both natural and human-derived

sources, this Assessment found the following:

(a) A substantial imbalance remains in estimates of source and sink magnitudes for both methyl bromide and

methyl chloride; known sinks outweigh sources for both of these gases. New sources of methyl bromide from

individual crops and ecosystems have been identified, and new sources of methyl chloride from tropical plants

have been discovered. These findings have narrowed the budget imbalances for both of these gases. Additional

studies continue to show that the ocean is a small (10-20 Gg yr-1) net sink for atmospheric methyl bromide,

SOURCE GASES

1.2
although this results from a balance of large production and loss terms. Our understanding of the saturation of

methyl bromide in ocean waters has been refined, but the estimates of net flux between the ocean and atmosphere remain essentially unchanged.

(b) Twentieth-century trends of methyl bromide and methyl chloride have been inferred for the Southern Hemisphere

from analyses of firn air. Provided these gases are neither produced nor destroyed in the firn, these trends suggest 20th century increases in the Southern Hemisphere of about 3 ppt for methyl bromide and 50 ppt for methyl

chloride.

(c) The best estimate for the global lifetime of methyl bromide remains at 0.7 (0.5-0.9) years. Additional studies

directly related to estimating loss processes for methyl bromide have narrowed the uncertainties slightly, but

they do not suggest large revisions to this lifetime. The fraction of emissions derived from industrially produced methyl bromide is unchanged at 10-40% based upon our current understanding of source and sink

magnitudes.

• Approaches to further accelerating the date of the recovery of the ozone layer are limited. This Assessment has

made hypothetical estimates of the upper limits of improvements that could be achieved if global anthropogenic

production of ozone-depleting substances were to stop in 2003 or if global anthropogenic emissions of ozonedepleting substances were to stop in 2003. Specifically:

Production. Relative to the current control measures (Beijing, 1999) and trends in recent production data, the

equivalent effective chlorine atmospheric loading above the 1980 level, integrated from 2002 until the 1980 level

is reattained (about 2050), could be decreased by the following amounts:

• 5%, if production of hydrochlorofluorocarbons (HCFCs) were to cease in 2003.

• 4%, if production of chlorofluorocarbons (CFCs) were to cease in 2003.

• 4%, if production of methyl bromide were to cease in 2003.

• 1%, if production of halons were to cease in 2003.

• 0.3%, if production of methyl chloroform were to cease in 2003.

Emissions. Similarly, the equivalent effective chlorine atmospheric loading above the 1980 level, integrated from

2002 until the 1980 level is reattained (about 2050), could be decreased by the following amounts:

• 11%, if emissions of halons were to cease in 2003.

• 9%, if emissions of chlorofluorocarbons (CFCs) were to cease in 2003.

• 9%, if emissions of hydrochlorofluorocarbons (HCFCs) were to cease in 2003.

• 4%, if emissions of methyl bromide were to cease in 2003.

• 3%, if emissions of carbon tetrachloride were to cease in 2003.

• 2%, if emissions of methyl chloroform were to cease in 2003.

The decreases calculated for reduced production and emissions scenarios would be about a factor of 2 smaller if the

decreases were compared with the loading integrated from 1980, which is when significant ozone depletion was

first detected. Furthermore, the decreases calculated for the integrated equivalent effective chlorine atmospheric

loading would be smaller if the cessation in production or emission occurred later than 2003.

The hypothetical elimination of all anthropogenic production of all ozone-depleting substances would advance the

return of stratospheric loading to the pre-1980 values by about 4 years. The hypothetical elimination of all emissions derived from anthropogenic production of all ozone-depleting substances would advance the return of stratospheric loading to the pre-1980 values by about 10 years.

SOURCE GASES

1.3

SOURCE GASES

1.5

1.1 INTRODUCTION

This chapter provides an update on scientific

progress since the previous Scientific Assessment of Ozone

Depletion (WMO, 1999) regarding ozone-depleting substances (ODSs) in the atmosphere. This includes a discussion of the latest available data for observed trends in

the troposphere and stratosphere, emissions, lifetimes,

Ozone Depletion Potentials (ODPs), and Global Warming

Potentials (GWPs). On the basis of this updated information, we relate observations of ODSs in the atmosphere to

expectations, and discuss the evidence that the Montreal

Protocol is effectively reducing ozone-depleting gases in

the atmosphere.

Also explored in this chapter are the potential future

mixing ratios of halocarbons based upon the current

Montreal Protocol and recent trends in halocarbon production data. Scenarios are presented to investigate a

range of possible future halocarbon mixing ratios. A

number of hypothetical “cases” are also explored to

demonstrate how production and emission of different

ODSs will affect atmospheric halogen burdens in the

future.

A discussion of non-ozone-depleting gases, such

as carbon dioxide and methane, is also included here.

Although they do not participate directly in ozonedestroying reactions, these gases influence stratospheric

ozone indirectly by affecting the availability of inorganic

chlorine and bromine, or by affecting stratospheric temperatures and thus the occurrence and persistence of polar

stratospheric clouds, among other effects.

The discussions here focus on all significant ODSs

with atmospheric lifetimes long enough that they are reasonably well mixed in the troposphere. This lifetime

cutoff was taken to be 0.5 years. Classic methods for calculating ODPs for gases with lifetimes longer than 0.5

years are assumed to be valid. The details regarding

budgets, sources, sinks, ODPs, etc., for gases with shorter

lifetimes are discussed in Chapter 2.

1.2 HALOGENATED OZONE-DEPLETING

GASES IN THE ATMOSPHERE

Atmospheric measurements of ozone-depleting

substances provide a foundation for understanding

changes in Earth’s protective ozone layer. Since the previous Assessment (WMO, 1999) additional measurements

have refined our understanding of the amounts, distributions, and changes in ozone-depleting gases in the atmosphere. For example, measurements suggest that inorganic

chlorine in the stratosphere has now stabilized. This

plateau arises as a result of the slow but continued overall

decline in organic chlorine observed globally in the troposphere since 1992-1994. Also, new results from the

analysis of firn air provide further evidence that most

ozone-depleting gases are entirely of human origin.

Details regarding these results and others are found here.

In this chapter and in subsequent ones, the terms

concentration, mixing ratio, volume mixing ratio

(assuming ideal gas behavior), abundance, amount, and

loading refer to dry air mole fraction. These mole fractions are expressed, for example, as parts per trillion (ppt;

pmol mol-1

). The term “ton” is used to represent a metric

ton, which is 106 grams. Also, the term “Montreal

Protocol” is used to indicate the fully amended and

adjusted Montreal Protocol as of 2002 (Beijing

Amendments), unless otherwise specified.

1.2.1 Updated Atmospheric Observations

of Ozone-Depleting Gases

1.2.1.1 CHLOROFLUOROCARBONS (CFCS)

Updated ground-based measurements of chlorofluorocarbons (CFCs) show continued increases for global

surface mixing ratios of CFC-12 (CCl2F2). Surface

mixing ratios of CFC-113 (CCl2FCClF2) peaked around

1996 and have been decreasing slowly thereafter. The

decrease noted for CFC-11 (CCl3F) mixing ratios in 1996

has continued (Figure 1-1; Table 1-1; Montzka et al., 1999;

Prinn et al., 2000). Observed rates of change in the lower

atmosphere for all three of these CFCs were slightly

smaller in 2000 than in 1996 (Table 1-1; Prinn and Zander

et al., 1999). Although emissions of CFC-12 declined

substantially during the 1990s, mixing ratios of CFC-12

continue to increase because emissions are still larger than

the small losses associated with its long atmospheric lifetime (100 years).

Calibration differences between three global

ground-based measurement networks—the Atmospheric

Lifetime Experiment/Global Atmospheric Gases Experiment/Advanced GAGE (ALE/GAGE/AGAGE; denoted

“AGAGE” in this chapter) network, the National Oceanic

and Atmospheric Administration/Climate Monitoring

and Diagnostics Laboratory (NOAA/CMDL; denoted

“CMDL” in this chapter) network, and the University of

California at Irvine (UCI) network—are on the order of

1-2% for CFC-11 and -12, and are slightly smaller for

CFC-113. The trends measured at Earth’s surface by these

three networks are similar in recent years (Table 1-1).

Updated column abundance measurements of CFC12 from the ground also show continued increases in the

atmospheric burden of this gas (Figure 1-2; Zander et al.,

2000). The mean rate of accumulation, however, has slowed
SOURCE GASES

1.6

CFC-114

5

1 0

1 5

ppt

2

4

6

8

ppt

1975 2000 2025 2050

year

CFC-115

CFC-13

CFC-114a

E0

Ab,Am,P0

E 0 Ab,Am,P0

CFC-113

20

40

60

80

ppt

975 2000 2025 2050

year

E0,P0

Am

A b

CFC-12

100

200

300

400

500

ppt

1975 2000 2025 2050

year

E0,P0

Am

Ab

CFC-11

50

100

150

200

250

ppt

975 2000 2025 2050

year

E 0

Ab,Am

P0

Figure 1-1. Past and potential future atmospheric mixing ratios of CFCs. Measurements of ambient air (solid

lines) and histories inferred from measurements of firn air (long-dashed lines) define past burdens at Earth’s

surface. Potential future mixing ratios (short-dashed lines) have been calculated for different scenarios that

are described in Section 1.8 (Ab, best guess; E0, zero emissions (not attainable, but defines a lower limit);

Am, maximum allowed production; P0, zero production). Data sources for measurements of ambient air:

AGAGE global monthly means (solid green lines: Prinn et al., 2000); CMDL global monthly means (solid red

lines: Montzka et al., 1999); UCI global quarterly means (solid purple lines: D.R. Blake et al., 1996); and

University of East Anglia (UEA) results from Cape Grim, Tasmania (for 41°S only) (solid blue lines: Oram,

1999). See the notes to Table 1-1 for more details regarding sampling frequencies and techniques. Antarctic

histories have been derived from firn air extracted at the South Pole (red dashed lines: Butler et al., 1999) and

Law Dome, Antarctica (green dashed lines: Sturrock et al., 2002). The shading shows the range of mixing

ratios encompassed by the future scenarios.
from 4.97 × 1013 molec cm−2 yr

−1 (or +0.71% yr−1

) in 1998

to 3.06 × 1013 molec cm−2 yr

−1 (or +0.43% yr−1

) in 2000.

Updated balloon-based measurements at northern

midlatitudes also reveal continued increases for CFC-12

in the stratosphere, up to 21 km altitude. Growth rates

between 1997 and 2001, however, are between 2 and 3

ppt yr

−1

, which is less (significant at the 2σ level) than

observed between 1978 and 1990, and also less (significant at the 1σ level) than observed between 1990 and 1997

(updated work from Engel et al., 1998).

Measured trends and atmospheric distributions

continue to suggest that CFC emissions in 2000 were substantially smaller than in the late 1980s and early 1990s

(Montzka et al., 1999; Prinn et al., 2000). The most recent

trends and interhemispheric gradients suggest, however,

that emissions are not yet insignificant. Persistent mixing

ratio gradients (or lack thereof) across latitudes on hemispheric scales can provide some indication of emission

locations, but they are generally not useful for delineating

emission rates from any individual country despite assertions to the contrary in some studies (Libo et al., 2001).

A number of less abundant CFCs have been measured in the atmosphere. CFC-13 (CClF3), CFC-114

(CClF2CClF2), and CFC-115 (CClF2CF3) have found use

as specialist refrigerants (R-13, R-114, R-115), aerosol

propellants (CFC-114, -115), and foam-blowing agents

(CFC-114). Small amounts of CFC-13 are also emitted

during aluminum production (Harnisch, 1997) and during

the manufacture of CFC-12. CFC-114a (CCl2FCF3),

which is an isomer of CFC-114 (CClF2CClF2), has also

been detected in the background atmosphere (Oram, 1999;

Culbertson et al., 2000). The most likely origin of CFC114a is as a byproduct of the CFC-114 production process

(Chen et al., 1994).

Measurements of air samples collected at Cape

Grim, Tasmania (41°S), show that the abundance of all

four of these minor CFCs increased substantially between

1978 and the mid-1990s (Figure 1-1; Table 1-1; Oram,

1999). During the early 1990s, however, the growth rates

began to decline and by 1995 had all fallen to below 0.4

ppt yr

−1

. In late 1995, mixing ratios at Cape Grim were

3.5 ppt (CFC-13), 16.5 ppt (CFC-114), 1.8 ppt (CFC114a), and 7.5 ppt (CFC-115) (Table 1-1).

More recent measurements from Cape Grim show

mid-2000 mixing ratios of about 8 ppt and 16.7 ppt for

CFC-115 and -114, respectively (Sturrock et al., 2001).

CFC-115 was increasing at 0.1 ppt yr

−1

, while CFC-114

had stopped growing altogether. Mangani et al. (2000)

similarly reported negligible growth for CFC-114 in

recent years from samples collected in Antarctica. For

these gases, direct comparisons between reported results

are difficult because they often do not overlap in time,

and, in the case of CFC-114, the AGAGE and Mangani et

al. (2000) measurements may include significant contributions from both C2Cl2F4 isomers (CFC-114 and -114a).

Culbertson et al. (2000) measured CFC-13 and

-114a in air samples collected at a rural, continental site

in the United States in March-April 2000. The mean

SOURCE GASES

1.7

CALENDAR YEAR

1986.0 1988.0 1990.0 1992.0 1994.0 1996.0 1998.0 2000.0 2002.0

C

OL

UMN

A

B

U

N

DA

NC

E (× 10

15 molec./cm2

) 1

2

5

6

7

8 CFC-12 and HCFC-22 above Jungfraujoch

CFC-12

HCFC-22

Pressure normalized monthly means

Pressure normalized June to November monthly means

Polynomial fit to filled datapoints

NPLS fit (20% ) Figure 1-2. The time evolution of the monthly mean

vertical column abundances

of CFC-12 and HCFC-22

above the Jungfraujoch station, Switzerland, from 1985

to 2002 (update of Zander et

al., 2000). Polynomial fits to

the June to November

columns are represented by

the solid lines; nonparametric least-square fits

(NPLS, dashed lines) are

also shown but are noticeable only for CFC-12.
SOURCE GASES

1.8

Table 1-1. Mixing ratios and growth rates of some important ozone-depleting substances.

Chemical Common or Mixing Ratio (ppt) Growth (1999-2000) Laboratory, Method

Formula Industrial Name 1996 1998 2000 (ppt yr–1) (% yr–1)

CFCs

CCl2F2 CFC-12 532.4 538.4 542.9 2.3 0.42 AGAGE, in situ a

525.9 531.0 534.5 1.9 0.35 CMDL, in situ a

523.2 529.3 534.0 1.8 0.34 CMDL, flasks a

526.2 532.0 535.7 1.9 0.35 UCI, flasks a

CCl3F CFC-11 265.6 263.0 260.5 –1.1 –0.41 AGAGE, in situ a

270.5 267.2 263.2 –2.0 –0.76 CMDL, in situ a

269.5 265.9 262.6 –1.5 –0.56 CMDL, flasks a

266.3 263.7 261.0 –1.0 –0.39 UCI, flasks a

CClF3 CFC-13 3.5 UEA, SH, flasks b

CCl2FCClF2 CFC-113 83.2 82.9 82.0 –0.35 –0.43 AGAGE, in situ a

84.2 83.0 82.1 –0.32 –0.39 CMDL, flasks b

83.1 82.1 81.1 –0.49 –0.60 UCI, flasks a

CClF2CClF2 CFC-114 16.5 UEA, SH, flasks b

CCl2FCF3 CFC-114a 1.8 UEA, SH, flasks b

CFC-114, -114a 17.1 17.2 –0.10 –0.58 AGAGE, in situ b

CClF2CF3 CFC-115 7.8 8.1 0.16 0.20 AGAGE, in situ b

7.5 UEA, SH, flasks b

Halons

CBrClF2 Halon-1211 3.9 4.1 0.13 3.2 AGAGE, in situ b

3.5 3.8 4.0 0.10 2.5 CMDL, flasks a, b

3.4 3.6 3.9 0.12 3.2 UCI, flasks a

3.8 4.2 4.4 0.13 2.9 UEA, SH, flasks b

CBrF3 Halon-1301 2.8 2.9 0.08 2.8 AGAGE, in situ b

2.3 2.5 2.6 0.06 2.4 CMDL, flasks a

2.0 2.2 2.3 0.05 2.2 UEA, SH, flasks b

CBrF2CBrF2 Halon-2402 0.48 CMDL, flasks b

0.42 0.42 0.43 0.001 0.2 UEA, SH, flasks b

CBr2F2 Halon-1202 0.037 0.044 UEA, SH, flasks b

Chlorocarbons—see also Chapter 2

CH3Cl Methyl chloride 538 536 AGAGE, SH, in situ b

588 OGI, flasks c

CCl4 Carbon 100.5 98.2 96.1 –0.94 –0.97 AGAGE, in situ a

tetrachloride 103.2 101.9 99.6 –0.95 –0.95 CMDL, in situ a

103.1 101.4 99.2 –1.03 –1.03 UCI, flasks a

CH3CCl3 Methyl 90.3 64.6 45.4 –8.7 –17 AGAGE, in situ a

chloroform 96.9 68.9 46.4 –10.2 –20 CMDL, in situ a

92.3 65.7 45.7 –9.1 –18 CMDL, flasks b

93.6 71.7 47.6 –13.0 –24 UCI, flasks a

HCFCs

CHClF2 HCFC-22 122.4 132.7 143.2 5.4 3.8 AGAGE, in situ a

121.5 131.4 141.9 5.1 3.7 CMDL, flasks b

CH3CCl2F HCFC-141b 9.5 13.0 1.8 15 AGAGE, in situ b

5.4 9.1 12.7 1.7 15 CMDL, flasks b

4.3 UT, flasks b
mixing ratio of CFC-13 was 3.6 ± 0.3 ppt, which is in reasonable agreement with the measurements of Oram

(1999). Conversely, the mixing ratio of CFC-114a was

7.9 ± 0.8 ppt, which is substantially higher than that

reported by Oram (1999).

1.2.1.2 HALONS

Although the phaseout was imposed on halon production in developed countries earlier than for all other

halocarbons, mixing ratios of both Halon-1211 (CBrClF2)

and Halon-1301 (CBrF3) continued to increase in the

atmosphere in 2000 (Figure 1-3). Global surface mixing

ratios of individual halons were less than 5 ppt in 2000

(Table 1-1). Since the previous Assessment (WMO, 1999)

a new set of regular measurements has become available

(Prinn et al., 2000; Sturrock et al., 2001). Among the laboratories studying long-term trends of halons in the atmosphere, calibration differences are about 10-15% (UCI <

CMDL < AGAGE < University of East Anglia (UEA)) for

Halon-1211 and about 25% (UEA < CMDL < AGAGE)

for Halon-1301 (Butler et al., 1998; Fraser et al., 1999;

Montzka et al., 1999).

For Halon-1211, all four laboratories report rates

of increase in 2000 of about 3% yr-1

, or 0.1 to 0.13 ppt

yr-1 (Table 1-1). For Halon-1301, CMDL, UEA, and

AGAGE data suggest rates of increase of 0.05-0.08 ppt

yr-1 or 2-3% yr-1 in 2000 (Table 1-1). For both these

halons, these rates of increase are only slightly slower

than observed during the first half of the 1990s.

Two less abundant halons have been further monitored in the atmosphere. Halon-2402 (CBrF2CBrF2) is

present at between 0.4 and 0.5 ppt in the global atmosphere (Table 1-1; Butler et al., 1998; Fraser et al., 1999),

and its rate of increase has slowed dramatically since the

mid-1990s (Fraser et al., 1999; Figure 1-3). Halon-1202

(CBr2F2) has been measured in background air at mixing

ratios of 0.04-0.05 ppt (Table 1-1; Figure 1-3; Engen et

al., 1999; Fraser et al., 1999). During 1995-1996 the abundance of Halon-1202 was increasing at 17% yr-1 (0.007

ppt yr-1) in the Southern Hemisphere, although updated

measurements by UEA suggest that this rate is much

smaller now (Figure 1-3). Although this halon has been

used by the military in a few minor applications, most is

produced from over-bromination during the production of

Halon-1211 (UNEP, 1998b).

1.2.1.3 CARBON TETRACHLORIDE (CCl4)

Global surface mixing ratios of carbon tetrachloride (CCl4) have decreased since about 1990; mixing ratios

in 2000 were between 95 and 100 ppt (Figure 1-4; Table

SOURCE GASES

1.9

Table 1-1, continued.

Chemical Common or Concentration (ppt) Growth (1999-2000) Laboratory, Method

Formula Industrial Name 1996 1998 2000 (ppt yr–1) (% yr–1)

CH3CClF2 HCFC-142b 10.4 12.5 1.1 9.4 AGAGE, in situ b

7.7 9.6 11.7 1.0 8.9 CMDL, flasks b

9.2 UT, flasks b

CHCl2F HCFC-21 0.29 UEA, SH, flasks b

CHCl2CF3 HCFC-123 0.03 UEA, SH, flasks b

CHClFCF3 HCFC-124 0.89 1.34 0.35 30 AGAGE, in situ b

Bromocarbons—see also Chapter 2

CH3Br Methyl bromide 9-10 Many d

8.4 8.1 AGAGE, SH, in situ b

Global mixing ratios and growth rates at Earth’s surface unless otherwise specified. AGAGE: 4-5 in situ electron capture detection (ECD) sampling

sites, 2 in situ gas chromatography-mass spectrometry (GC-MS) sampling sites, 2 flask sampling sites; CMDL: 5 in situ sampling sites, 8-10 remote

flask sampling sites; UEA: University of East Anglia, analysis of archive flasks and regularly sampled flasks filled at Cape Grim, Australia (41°S);

UCI: quarterly sampling from a multitude of sites between 47°S and 71°N.; UT: University of Tokyo, 2 flask sampling sites; OGI: Oregon Graduate

Institute, quarterly flask sampling at ~40 sites. NH, Northern Hemisphere; SH, Southern Hemisphere.

Data sources: AGAGE: Prinn et al. (2000); Sturrock et al. (2001); ftp://cdiac.esd.ornl.gov/pub/ale_gage_Agage. CMDL: Butler et al. (1999); Montzka

et al. (1999, 2000); Hall et al. (2002); http://www.cmdl.noaa.gov. UEA: Fraser et al. (1999); Oram et al. (1995); Oram (1999). UCI: D.R. Blake

et al. (1996); N.J. Blake et al. (2001). UT: Shirai and Makide (1998).

a Measurements by gas chromatography with electron capture detection.

b Measurements by gas chromatography with mass spectrometry detection. c Results from OGI for CH3Cl have been scaled by 550/600; see Kurylo and Rodríguez et al. (1999).

d See Section 1.5.
SOURCE GASES

1.10

0.1

0.2

0.3

0.4

ppt

1975 2000 2025 2050

year

Ab,Am,P0

E 0 0.01

0.02

0.03

0.04

0.05

ppt

1975 2000 2025 2050

year

E0,P0

A b A m

2

3

ppt

1975 2000 2025 2050

year

E 0

P0

Ab

Am

2

3

4

ppt

1975 2000 2025 2050

year

P0

A m

E 0

A b

Halon-2402

Halon-1211 Halon-1301

Halon-1202

Figure 1-3. Past and potential future atmospheric mixing ratios of halons. Measurements of ambient air

(solid lines) and histories inferred from measurements of firn air (long-dashed lines) define past burdens at

Earth’s surface. Potential future mixing ratios (short-dashed lines) have been calculated for different scenarios that are described in Section 1.8 and the caption to Figure 1-1. Data inferred from firn air are from the

same sources as described in Figure 1-1. Data sources for measurements of ambient air: UEA from Cape

Grim, Tasmania (for 41°S only) (solid blue lines: Fraser et al., 1999); AGAGE global mean data (solid green

lines: Sturrock et al., 2001); UCI global quarterly means (solid purple lines: Blake et al., 2001); and CMDL

global means (solid red lines: Butler et al., 1998; Montzka et al., 1999). The shading shows the range of

mixing ratios encompassed by the future scenarios. See the notes to Table 1-1 for more details regarding

sampling frequencies and techniques.
1-1; Simmonds et al., 1998a; Montzka et al., 1999; Prinn

et al., 2000). The observed rates of change have remained

fairly constant since 1993 at about –1% yr-1 or –1 ppt

yr-1. The interhemispheric difference has also been fairly

constant at about 2% (North > South) since 1993 and suggests that significant emissions of carbon tetrachloride

remain (see Section 1.6). Calibration differences between

CMDL, UCI, and AGAGE are 3 to 4%.

1.2.1.4 METHYL CHLOROFORM (CH3CCl3)

The rapid decline in emissions of methyl chloroform (CH3CCl3) and its relatively short lifetime have

together resulted in rapidly decreasing mixing ratios in

recent years (Figure 1-4; Montzka et al., 2000; Prinn et

al., 2001). The global mean surface mixing ratio in 2000

was approximately 46 ppt (Table 1-1), compared with the

maximum of 130 ppt observed in 1992. Hemispheric differences were 2 to 3% in 2000, or much smaller than in

earlier years.

Fairly constant exponential decay with a time constant of (5.5 yr)-1 was observed for methyl chloroform

during 1998-2000 (Montzka et al., 2000). This implies

that the absolute rate of decline for methyl chloroform is

becoming smaller: it peaked at –14 to –15 ppt yr-1 in

1995-1996, and was one-third less in 1999-2000, or about

–10 ppt yr-1 (Table 1-1).

Scale differences between CMDL and AGAGE

were stated as being ~10% in Prinn and Zander et al.

(1999), and even larger among a broader range of laboratories. Since then scale revisions by AGAGE (Scripps

Institution of Oceanography (SIO), SIO-93 to SIO-98;

Prinn et al., 2000) and CMDL (Hall et al., 2002) suggest

that the AGAGE-CMDL differences are now on the order

of <3%. There is a small time dependence to this difference. Similar mixing ratios are reported by UCI in 2000

(Table 1-1).

1.2.1.5 HYDROCHLOROFLUOROCARBONS (HCFCS)

Updated measurements of hydrochlorofluorocarbons (HCFCs) indicate that global mixing ratios of the

three most abundant HCFCs continue to increase in the

atmosphere, owing to sustained emissions (Figure 1-5).

Updated measurements suggest global mixing ratios for

HCFC-22 (CHClF2) of 140-145 ppt in 2000 and a fairly

constant growth rate of about 5 ppt yr-1 (Table 1-1; Figure

1-5; Simmonds et al., 1998b; Montzka et al., 1999).

Sturrock et al. (2001) reported slightly faster increases

from samples collected only at Cape Grim, Tasmania (6.3

ppt yr-1). Hemispheric differences have been fairly constant in recent years at about 10% (North > South).

SOURCE GASES

1.11

3 0

6 0

9 0

120

ppt

1975 2000 2025 2050

year

E 0

P 0 Ab,Am

20

40

60

80

100

ppt

1950 1975 2000 2025 2050

year

E0

Ab,P0

A m

carbon tetrachloride

m ethyl chloroform

Figure 1-4. Past and potential future atmospheric

mixing ratios of carbon tetrachloride (CCl4) and methyl

chloroform (CH3CCl3). Measurements of ambient air

(solid lines) and histories inferred from measurements

of firn air (long-dashed lines) define past burdens at

Earth’s surface. Potential future mixing ratios (shortdashed lines) have been calculated for different

scenarios that are described in Section 1.8 and the

caption to Figure 1-1. Data sources and shading are

the same as described in Figure 1-1.
Mean total column abundances of HCFC-22 measured above the Jungfraujoch station, Switzerland, in 1995

were 1.66 ¥ 1015 molec cm-2

, increasing by 8.71 ¥ 1013

molec cm-2 yr-1 (or 5.2% yr-1); for 2000, these figures

were 2.11 ¥ 1015 molec cm-2

, increasing by 8.93 ¥ 1013

molec cm-2 yr-1 (or 4.2% yr-1), thus indicating that the

absolute rate of atmospheric accumulation during those 5

years did not change substantially (Figure 1-2; an update

of Zander et al., 2000).

Updated measurements show that the global mean

surface mixing ratio of HCFC-141b (CH3CCl2F) rose

from about 5 ppt in mid-1996 to nearly 13 ppt in mid-2000

(Table 1-1; Figure 1-5; Montzka et al., 1994, 1999;

Simmonds et al., 1998b; Prinn et al., 2000; Sturrock et al.,

2001). The current growth rate is slightly less than 2 ppt

yr-1

. Northern Hemispheric mean mixing ratios were

approximately 3 ppt higher than Southern Hemispheric

means in mid-2000.

SOURCE GASES

1.12

5

1 0

1 5

2 0

2 5

3 0

ppt

1975 2000 2025 2050

year

E 0

P0

A m

A b

10

20

30

40

ppt

975 2000 2025 2050

year

E 0

P0

A b

A m

40

80

120

160

ppt

975 2000 2025 2050

year

E 0

A b P 0

Am

HCFC-22

HCFC-141b

HCFC-142b

Figure 1-5. Past and potential future atmospheric

mixing ratios of HCFCs. Measurements of ambient air

(solid lines) and histories inferred from measurements

of firn air (long-dashed lines) define past burdens at

Earth’s surface. Potential future mixing ratios (shortdashed lines) have been calculated for different scenarios that are described in Section 1.8 and the caption to Figure 1-1. Data sources: mixing ratios from

Antarctic firn (long-dashed green line: Sturrock et al.,

2002); AGAGE global means (solid green lines: Prinn

et al., 2000; Sturrock et al., 2001); CMDL global means

(solid red lines: Montzka et al., 1999); UEA from Cape

Grim, Tasmania (for 41°S only) (solid blue lines: Oram

et al., 1995); SIO from Cape Grim (41°S only) (solid

black line: Miller et al., 1998). See the notes to Table

1-1 for more details regarding sampling frequencies

and techniques. The shading shows the range of

mixing ratios encompassed by the future scenarios.
Global mean surface mixing ratios of HCFC-142b

(CH3CClF2) have risen from 7 to 8 ppt in mid-1996 to

about 12 ppt in mid-2000, with a current growth rate of

1.0 ppt yr-1 (Table 1-1; Figure 1-5; Montzka et al., 1994;

1999; Simmonds et al., 1998b; Prinn et al., 2000; Sturrock

et al., 2001). Northern Hemispheric means were 2 to 2.5

ppt higher than Southern Hemispheric means in mid-2000.

Calibration differences are about 5% or less for HCFC22, -141b, and -142b among laboratories reporting results

recently (Table 1-1).

Since the previous Assessment, three other HCFCs

have been reported in the background atmosphere: HCFC123 (CHCl2CF3), HCFC-124 (CHClFCF3), and HCFC-21

(CHCl2F) (Table 1-1). From measurements at Cape Grim,

Sturrock et al. (2001) reported late-1998 tropospheric

mixing ratios of 0.53 ppt and 0.1 ppt for HCFC-124 and

-123, respectively, the former having a growth rate of 0.06

ppt yr-1 (13%). Oram (1999) found similarly low levels

of HCFC-123 at Cape Grim from archived samples collected between 1978 and 1993, but noted that the mixing

ratio doubled between 1990 and 1993. Oram (1999) also

reported measurements of HCFC-21 over the same period.

Mixing ratios were typically in the range 0.2-0.4 ppt, with

no significant trend.

1.2.1.6 METHYL BROMIDE (CH3Br)

Our understanding of the concentration of methyl

bromide (CH3Br) in the atmosphere during the 1990s

remains much the same as was given in the previous

(1998) Assessment (WMO, 1999). No new global trends

have been reported since those of Khalil et al. (1993).

Recent data from research cruises encompassing a wide

latitudinal span in both hemispheres (Yokouchi et al.,

2000a; Li et al., 2001; Groszko and Moore, 1998; King et

al., 2000), from aircraft missions (Schauffler et al., 1999),

and from two midlatitude, ground-based sites (Miller,

1998) suggest that the mean global mixing ratio of atmospheric CH3Br ranged from 9 to 10 ppt before reductions

in industrial production began in the late 1990s, consistent with that reported in the 1998 Assessment. These

new studies, along with that of Wingenter et al. (1998),

suggest that the best estimate of the global, area-weighted,

mean hemispheric ratio lies between 1.2 and 1.3 (North >

South). This hemispheric ratio varies seasonally from

1.1 to 1.4, driven mainly by seasonality in Northern

Hemispheric mixing ratios. Recent published comparisons between the National Center for Atmospheric

Research (NCAR) and UCI show measurements agreeing

within an average of 7% (Schauffler et al., 1999).

The vertical distributions and weak tropospheric

gradient observed by N.J. Blake et al. (1996; 1997) are

supported by measurements of Schauffler et al. (1999).

The tropospheric gradient in mixing ratio of 3-14%

reported by Schauffler et al. (1999) is consistent with that

reported in the 1998 Assessment, which noted a 0-15%

vertical gradient in the tropospheric mixing ratio.

Additionally Schauffler et al. (1999) noted that CH3Br

mixing ratios in the lower stratosphere (20 km) were

reduced by 40% in the tropics and by 70% at Northern

Hemispheric midlatitudes (July) relative to amounts

observed in the troposphere.

1.2.1.7 METHYL CHLORIDE (CH3Cl)

Khalil and Rasmussen (1999) recently published a

detailed analysis of global measurements of methyl chloride (CH3Cl) conducted over 16 years (1981-1997) at locations distributed throughout both hemispheres, during

which methyl chloride reportedly decreased by about 4%.

As noted in Chapter 2 of the 1998 Assessment (Kurylo

and Rodríguez et al., 1999), Khalil and Rasmussen’s

global mean mixing ratio of 606 ppt is somewhat higher

than the average of 550 ± 30 ppt obtained by other investigators. New results continue to suggest global means

closer to 550 ppt than 600 ppt (Sturrock et al., 2001;

Yokouchi et al., 2000b; Table 1-1). Khalil and Rasmussen

(1999) reported a mean annual latitudinal distribution in

the CH3Cl mixing ratio in which the tropics were about

40 ppt higher than the poles, and they inferred that this

had to be caused by a tropical terrestrial source. They also

reported a seasonal cycle with an amplitude of about 10%,

explicable mainly in terms of seasonal changes in

hydroxyl radical (OH) abundance. In more recent work

involving land-based and shipboard measurements

extending from the Canadian Arctic through the Pacific

and Indian Oceans to the Antarctic between 1996 and

1998, Yokouchi et al. (2000b) found CH3Cl mixing ratios

ranging from 500 ppt in the high Arctic and Antarctic to

570 ppt near the equator. Although this latitudinal distribution was measured only over a fraction of a year, it is

similar to the annual mean distribution reported by Khalil

and Rasmussen (1999).

1.2.2 Twentieth Century Atmospheric

Histories for Halocarbons

from Firn Air

Since the previous Assessment (WMO, 1999), 20th

century atmospheric histories for halocarbons have been

reconstructed from the analysis of air trapped in snow

above glaciers, also known as firn air (Figures 1-1, 1-3,

1-4, and 1-5; Butler et al., 1999; Sturges et al., 2001a;

Sturrock et al., 2002). Conclusions drawn from the firnair results rely on the assumption that these halocarbons

are neither produced nor destroyed in the firn. These firnSOURCE GASES

1.13
air data show that mixing ratios of CFCs, halons, and

HCFCs in the oldest air sampled are generally less than

2% of the amounts measured today in the background

atmosphere. Furthermore, and consistent with these halocarbons not being destroyed or produced in firn air over

time, the 20th century atmospheric histories reconstructed

from firn-air results and firn diffusion models are reasonably consistent with calculated histories based on records

of industrial halocarbon production (Prinn et al., 2000).

Somewhat higher amounts of carbon tetrachloride (5 ppt)

and methyl chloroform (2-4 ppt) have been reported in

the deepest and oldest samples, although these amounts

were close to the detection limit of the instruments used

in these analyses. These results suggest that nonindustrial sources of these halocarbons are insignificant.

Mixing ratios for methyl chloride and methyl bromide were only 10-30% lower than present day in firn air

from Antarctica dating back to the early 1900s (Butler et

al., 1999; Sturges et al., 2001a). The results are consistent with the presence of substantial nonindustrial emissions of these gases. Consistent results were observed for

methyl bromide in firn samples at four different locations

in Antarctica. These suggest that levels of methyl bromide in the Southern Hemisphere have increased by about

3 ppt since the early 20th century and about 2 ppt since the

mid-20th century (see Section 1.5). Unfortunately,

Northern Hemispheric firn-air profiles of methyl bromide

show significant anomalies near the snow-ice transition, a

feature not observed in the Antarctic profiles. For methyl

chloride, firn data from both hemispheres suggest that the

atmospheric burden over the last half of the 20th century

increased by about 10%. Additional details and discussion of these results appear in Section 1.5.

1.2.3 Total Atmospheric Chlorine

1.2.3.1 TOTAL ORGANIC CHLORINE IN THE

TROPOSPHERE

Total organic chlorine (CCly) contained in longlived chlorine-bearing source gases continues to decrease

slowly in the lower atmosphere. In mid-2000, CCly was

~3.5 parts per billion (ppb) (Table 1-2), or about 5% lower

than the peak observed in 1992-1994. The decrease still

results primarily from the exponential decline observed

for methyl chloroform (Table 1-2; Table 1-1; Figure 1-4).

This situation is changing, however, because the absolute

rate of decrease observed in 1999-2000 for methyl chloroform had diminished by one-third compared with 1995-

1996 and will continue to lessen in the future (Montzka et

al., 1999; see also Section 1.8).

By 2000 total chlorine from aggregated CFCs was

no longer increasing (Table 1-2). Continued increases in

chlorine from CFC-12 (3.6-4.6 ppt Cl yr-1) in 2000 were

similar in magnitude to declines in chlorine from the sum

of CFC-11 and CFC-113 (Table 1-1).

The rate of increase in chlorine from HCFCs has

been fairly constant since 1996 (Tables 1-1 and 1-2).

HCFC-22 accounts for approximately 80% of chlorine

from HCFCs in today’s atmosphere and for about half of

the annual increase in chlorine from all HCFCs (5 ppt Cl

yr-1 out of about 10 ppt Cl yr-1). Continued increases in

SOURCE GASES

1.14

Table 1-2. Contributions of halocarbons to total organic chlorine (CCly) in the troposphere.

Total CCly (ppt Cl) Contribution to Total Rate of Change in Total

CCly (%) CCly (ppt Cl yr–1)

1996 2000 1996 2000 1996 2000

Aggregate CFCs 2161 2156 59 61 7.1 –0.6

CH3Cl 550 550 15 16 0.0 0.0 a

CCl4 407 384 11 11 –4.0 –3.8

Aggregate HCFCs 142 182 4 5 10.3 9.7

CH3CCl3 278 139 8 4 –40.8 –27.2

Short-lived gases b 100 100 3 3 0.0 a 0.0 a

Halon-1211 4 4 0 0 0.2 0.1

Total CCly 3642 3516 –27 –22

(–0.8% yr–1) (–0.6% yr–1)

Bold-faced type is used to highlight the largest changes from 1996 to 2000. Some differences for 1996 from WMO (1999) arise because of small recent

adjustments to absolute calibration scales and because the results presented here are an average of AGAGE and CMDL global means. Similar conclusions could be drawn with data from UCI (Table 1-1).

a Presumed to be zero; not well documented.

b Gases such as CH2Cl2, CHCl3, and C2Cl4.
HCFC-141b and -142b mixing ratios account for the additional annual chlorine increase.

The four minor CFCs accounted for about 50 ppt

of Cl, or less than 1.5% of the total tropospheric chlorine

burden in 2000 (CFC-13, -114, -114a, and -115). About

70% of this 50 ppt is from CFC-114. The current rate of

growth for chlorine in these minor CFCs is small (<1 ppt

Cl yr-1; Table 1-1).

1.2.3.2 TOTAL INORGANIC CHLORINE IN THE

STRATOSPHERE

In a stratosphere unperturbed by polar stratospheric

clouds or recent volcanism, more than 95% of inorganic

chlorine (Cly) is accounted for by the two reservoirs

hydrogen chloride (HCl) and chlorine nitrate (ClONO2)

(Zander et al., 1992, 1996). Furthermore, because tropospheric HCl and ClONO2 mixing ratios are small, the sum

of their vertical column abundances is a good surrogate

of the Cly loading and its evolution in the stratosphere.

Past reports of total column abundances of HCl and

ClONO2, as well as their sum (Cly) derived from observations at Jungfraujoch (Switzerland, 46.5°N) through 1999,

showed increases commensurate with trends in CCly at

Earth’s surface but delayed by 3-4 years owing to air transport rates between the troposphere and the stratosphere

(Prinn and Zander et al., 1999; Mahieu et al., 2000).

An update of measurements from the Jungfraujoch

now shows that the monthly mean total column abundances of HCl and ClONO2 have stopped increasing

(Figure 1-6). Updates of column abundance measurements of HCl at Kitt Peak (Arizona, U.S., 31.9°N)

(Rinsland et al., 1991; Wallace et al., 1997) and of HCl

and ClONO2 at Lauder (New Zealand, 45°S) (Matthews

et al., 1989; Reisinger et al., 1995) also show similar

overall long-term trends. These results provide robust

evidence that the loading of inorganic Cly in the unperturbed stratosphere has recently stabilized in response to

the production regulations on ozone-depleting substances

outlined in the Montreal Protocol. Within the uncertainty

of these measurements, the trends in total column HCl

and ClONO2 are consistent with the trends for chlorinated

organic trace gases measured at Earth’s surface.

1.2.3.3 HYDROGEN CHLORIDE (HCl) AT 55 KM

ALTITUDE IN THE STRATOSPHERE

An independent assessment of total Cly near the

stratopause has been provided by measurements of HCl

made since 1991 at 55 km altitude by the Halogen

Occultation Experiment (HALOE) instrument aboard the

Upper Atmosphere Research Satellite (UARS) (Russell et

al., 1996; Prinn and Zander et al., 1999). At that altitude,

most Cly (~ 93-95%, globally) is in the form of HCl. The

broad changes in mean global HCl at 55 km measured by

HALOE show a period of monotonically increasing HCl

mixing ratios before 1997, and much slower mean changes

thereafter (updates to Anderson et al., 2000; Figure 1-7).

These long-term trends are consistent with expectations

from ground-based measurements after considering lag

times associated with transport of tropospheric air to the

stratosphere and stratospheric mixing processes (Waugh

et al., 2001; Hall and Plumb, 1994).

What remain unexplained in the HALOE HCl data

are abrupt changes that occurred over shorter periods, such

as those observed in early 1997 (Waugh et al., 2001; Engel

et al., 2002) and in 2000 (Figure 1-7). These changes

strongly suggest that additional atmospheric processes

affect trends in HCl at 55 km (Randel et al., 1999;

Considine et al., 1999). Several possible mechanisms

likely to cause the early-1997, abrupt peaking of Cly at 55

km altitude were considered by Waugh et al. (2001) (such

as changes in transport and chemistry partitioning), but

none could explain a sharp decrease as early as 1997 in

upper stratospheric Cly, given the measured tropospheric

halocarbon trends and our understanding of atmospheric

transport rates and mixing processes. HALOE measurements of methane, however, are anticorrelated with HCl

since 1997. This suggests that transport may contribute

to the variability observed in the HCl record, but transport effects alone do not resolve these tendencies (personal communication, J. Russell III, Hampton University,

U.S., 2001).

1.2.3.4 STRATOSPHERIC CHLORINE COMPARED WITH

TROPOSPHERIC CHLORINE

Evaluations of total chlorine (Cltot) in the stratosphere have been derived from measurements of CCly and

Cly throughout the stratosphere with the Atmospheric

Trace Molecule Spectroscopy (ATMOS) shuttle-based

Fourier transform infrared (FTIR) spectrometer (Gunson

et al., 1996). Average stratospheric Cltot was measured at

2.58 ± 0.10 ppb in 1985 (Zander et al., 1992) and 3.53 ±

0.10 ppb in 1994 (Zander et al., 1996). From measurements of a suite of chlorine-bearing source-, sink-, and

reservoir species with a balloonborne FTIR spectrometer

launched on 8 May and on 8 July 1997 from Fairbanks

(Alaska, 64.8°N), Sen et al. (1999) derived a nearly constant value of Cltot, equal to 3.7 ± 0.2 ppb between 9 and

38 km altitude. Each of these studies confirmed that the

bulk of inorganic stratospheric chlorine is well explained

by the photodissociation of organic source gases measured in the troposphere after considering time lags associated with air transport and mixing processes.

SOURCE GASES

1.15
1.2.4 Total Atmospheric Bromine

1.2.4.1 TOTAL ORGANIC BROMINE IN THE

TROPOSPHERE

Total organic bromine (CBry) in the troposphere

results primarily from surface emissions of methyl bromide, halons, and short-lived bromocarbons. Recent

trends in CBry are uncertain because global atmospheric

trends in methyl bromide since 1992 are not yet well documented. Firn-air measurements suggest significant

increases for this gas during 1900-1998, but they do not

tightly constrain recent atmospheric trends. Changes in

the growth rate of methyl bromide are perhaps likely

during the 1990s because of the restrictions placed on

industrial production in developed countries then (see the

fully revised and amended Montreal Protocol).

Total organic bromine from halons has continued

to increase since 1996 owing to continued anthropogenic

production and use. The aggregate contribution of halons

SOURCE GASES

1.16

1983.0 1986. 0 1989.0 1992.0 1995.0 1998.0 2001.0

T

otal C

olu

m

n

A

b

u

n

d

a

nce (× 10

15 m

olec./cm2

) 0

1

2

3

4

5

6

Year

1983.0 1986.0 1989.0 1992.0 1995.0 1998.0 2001.0

2

3

4

Lagged by 3.5 yr

Measured

C

Cl

y (p

p

b)

HCl

ClONO2

Cl

y

June to November monthly means

Nonparametric LS fit (20%)

Inorganic chlorine above

Jungfraujoch

Global organic chlorine

at the Earth surface

Figure 1-6. Upper frame: Time series of June to November (to avoid significant variability during winter and

spring periods) monthly mean vertical column abundances of HCl (open circles) and ClONO2 (open triangles)

derived from solar observations at the Jungfraujoch between 1983 and 2001 (update from Mahieu et al.,

2000). Inorganic chlorine (Cl

y, filled triangles) is obtained by summing the HCl and ClONO2 column measurements. Lines represent nonparametric least-squares fits with 20% weighting. Lower frame: The temporal

evolution of global organic chlorine loading (CCl

y) determined from in situ measurements at the ground (solid

line: Prinn et al., 2000), and CCl

y lagged by 3.5 years (dashed line).
to CBry was 0.7 ppt in 1978, 6.7 ppt in 1996, and 7.7 ppt

in 2000, and was increasing at 0.2 ppt Br yr-1 in 2000

(Table 1-1; Figure 1-8).

Shorter-lived gases with predominantly natural

sources, such as bromoform (CHBr3) and dibromomethane (CH2Br2), also add to the bromine burden of

the atmosphere. Atmospheric measurements indicate

between 1 and 3 ppt of organic bromine is present near

the tropopause in the form of these gases (Shauffler et al.,

1998; Pfeilsticker et al., 2000). Although the precise

amount of organic plus inorganic bromine delivered to

the stratosphere by these short-lived gases is uncertain

(see Chapter 2), we estimate that in 2000 methyl bromide

accounted for about 50% of CBry, and the halons

accounted for about 40% of CBry.

1.2.4.2 TOTAL INORGANIC BROMINE IN THE

STRATOSPHERE

The burden and trend of inorganic bromine (Bry)

in the stratosphere can be estimated indirectly from measurements of organic source gases in the stratosphere

(Schauffler et al., 1998; Wamsley et al., 1998) and from

stratospheric bromine monoxide (BrO) measurements and

calculated BrO/Bry partitioning (Harder et al., 2000;

Pfeilsticker et al., 2000; Sinnhuber et al., 2002). In stratospheric air in which most organic source gases had

become oxidized, Bry was estimated in early 1999 to be

18.4 (+1.8/-1.5) ppt from organic precursor measurements and 21.5 ± 3.0 ppt from coincident BrO measurements and photochemical modeling (Pfeilsticker et al.,

SOURCE GASES

1.17

Figure 1-7. The time evolution of monthly mean HCl

at altitudes of 54-56 km between 70°N and 70°S as

measured by HALOE (green line: an update of

Anderson et al., 2000). From this HCl record, total

inorganic chlorine (Cly) mixing ratios at 55 km (red

line) and their ±1 standard deviations (orange

shading) were derived using weighted contributions

at 55 km altitude from all important source gases as

determined from NCAR 2-D model calculations

(Brasseur et al., 1990). Also shown are (i) CCly at

Earth’s surface based on the Ab scenario in

Madronich and Velders et al. (1999) (dashed blue

line), (ii) this surface result lagged by 6 years (solid

blue line), and (iii) this lagged CCly convolved with a

3-yr-wide “age spectrum” (black line: an update of

Waugh et al., 2001).

6

8

10

12

14

16

18

20

22

24

26

methyl

bromide

+ halons Bromine (ppt)

Year

methyl

bromide

Total

Inorganic

Figure 1-8. Past measured trends for bromine in

the troposphere (lines) and stratosphere (points):

global tropospheric bromine from methyl bromide

as measured in ambient air and firn air (short dashed

line: Khalil et al., 1993; Butler et al., 1999; Sturges

et al., 2001a); global tropospheric bromine from the

sum of methyl bromide plus halons as measured in

ambient air, archived air, and firn air (solid line: Butler

et al., 1999; Fraser et al., 1999; Sturges et al.,

2001a), or as derived from firn-air and global flask

samples (long-dashed line: Butler et al., 1998); and

total inorganic bromine derived from stratospheric

measurements of BrO and photochemical modeling

that accounts for BrO/Bry partitioning (open squares:

Harder et al., 2000; Fitzenberger et al., 2000;

Pfeilsticker et al., 2000). The years indicated on the

abscissa are sampling times for tropospheric data.

For stratospheric data, the date corresponds to

when that air was last in the troposphere (i.e., sampling date minus mean time in stratosphere).
2000; measurements were made at 25 km in air inferred

to have a 5.6-yr mean age). Lower in the stratosphere,

Bry is a much smaller fraction of total bromine; these same

investigators estimated Bry to be 1.5 ppt in air just above

the local Arctic tropopause (about 9.5 km).

1.2.4.3 STRATOSPHERIC BROMINE COMPARED WITH

TROPOSPHERIC BROMINE

Spectrometric BrO measurements allow for estimates of total stratospheric Bry. These measurements suggest an increase in Bry over time and a stratospheric

burden that is 4 to 6 ppt higher than indicated by the tropospheric burdens of halons and methyl bromide alone

(Figure 1-8; an update of Pfeilsticker et al., 2000).

Measurements of long- and short-lived gases at the

tropopause and throughout the stratosphere suggest that

an additional 1 to 3 ppt of bromine reaches the stratosphere in the form of organic short-lived gases, such as

dibromomethane and bromoform (Schauffler et al., 1998;

Wamsley et al., 1998; Pfeilsticker et al., 2000). This additional bromine does not appear to account for all the BrO

measured, however. The remaining difference may suggest the presence of unmeasured, brominated organic

gases reaching the stratosphere, or an influx of inorganic

bromine into the stratosphere (see Chapter 2), or calibration errors either in the measurement of BrO or the organic

source gases (Fitzenberger et al., 2000).

1.2.5 Effective Equivalent Chlorine

and Effective Equivalent

Stratospheric Chlorine

The net effect of changes in both atmospheric chlorine and bromine can be gauged roughly by computing

effective equivalent chlorine (EECl) and effective equivalent stratospheric chlorine (EESC) from ground-based

measurements of halocarbons (Prather and Watson, 1990;

Daniel et al., 1995; Montzka et al., 1996a, 1999; WMO,

1999). EECl is calculated with ground-based halocarbon

measurements, consideration of the enhanced efficiency

of bromine to deplete stratospheric ozone (a factor of 45

used here; see discussion of a in Section 1.4.4), and consideration of the relative rates at which halocarbons

decompose and release their halogen into the stratosphere.

Ground-based measurement networks of the most abundant ODSs indicate that total ozone-depleting halogen

measured in the troposphere peaked in 1992-1994 (Figure

1-9; Montzka et al., 1996a; Cunnold et al., 1997). Updated

observations show that EECl in tropospheric air declined

from 1995 through 2000 at a mean rate of about 1.2%

yr-1

, or 24 ppt EECl yr-1. As of mid-2000, EECl was about

5% below the peak that was observed in 1992-1994.

SOURCE GASES

1.18

-2.0%

-1.0%

0.0%

1.0%

2.0%

1990 1995 2000 2005 2010

Sample date Rate of Change (per yr)(b)

1850

1950

2050

2150

2250

1990 1995 2000 2005 2010

Sample date Effective Equivalent Cl (ppt) (a)

Figure 1-9. Burden and trends in aggregate tropospheric Cl plus Br from purely anthropogenic halocarbons expressed as effective equivalent chlorine.

(a) Effective equivalent chlorine estimated from

ground-based measurements of the major CFCs,

methyl chloroform, carbon tetrachloride, Halon1211, Halon-1301, and HCFCs (blue line with diamonds: Montzka et al., 1999; red line with squares:

Prinn et al., 2000), and from the same gases as in

scenario Ab (thin smooth line). The time scale on

the x-axis refers to the date tropospheric air was

sampled; to convert to effective equivalent stratospheric chlorine (EESC), add approximately 3 years

to these dates. (b) The rate of change in effective

equivalent chlorine as measured by [AGAGE +

CMDL halon] measurements (red squares); CMDL

measurements only (blue diamonds); [AGAGE +

UEA halon] measurements (green plus symbols:

halon data from Fraser et al., 1999); and in scenario

Ab (thin solid line). Rates of change correspond to

the 12-month difference over the previous 12

months. In both panels, total bromine was multiplied

by 45, and the absolute fractional release assumed

for CFC-11 was 0.8 (see Section 1.4.4).
Smaller declines and slower rates of decline are inferred

when halon measurements by UEA at Cape Grim are considered (Figure 1-9; update of Fraser et al., 1999).

Although decreases observed for methyl chloroform are the primary reason for the turnover in EECl, the

influence of methyl chloroform on EECl is diminishing.

Furthermore, the increases still observed for CFC-12 and

halons are slowing the decline of EECl (Fraser et al., 1999;

Montzka et al., 1999). As a result of these influences,

EECl decreased in 2000 at a rate that was only about twothirds of the rate 3 to 4 years earlier (Figure 1-9). As the

influence of methyl chloroform continues to lessen, a sustained decrease in EECl is assured only if trends for other

halocarbons decrease further or become more negative

(Montzka et al., 1999; see also Section 1.8).

Whereas EECl reflects the time evolution of equivalent effective halogen only in the troposphere, EESC provides an estimate of that burden in the stratosphere. EESC

is derived by simply adding a 3-year time lag to EECl (see

Section 1.8). Although EESC provides a useful measure

to gauge the combined influence of tropospheric chlorine

and bromine trends on stratospheric ozone-depleting

halogen, it has important limitations. First, stratospheric

halocarbon abundances lag tropospheric abundances by a

few months to 6 years, depending upon location. Despite

this range, a lag of 3 years is generally used in the calculation of EESC to approximate the mean transport time

for air from the surface to the lower-mid stratosphere,

which is where most ozone depletion has been observed.

The 3 years is typical of the average age of air (time since

air was in the troposphere; see Section 1.4.3) in these

regions. Second, stratospheric halocarbon burdens are

not determined by tropospheric burdens at a single time

in the past. They instead represent a mean of tropospheric

burdens from a range of earlier years (Hall and Plumb,

1994); this can result in errors with this simple technique

especially when atmospheric mixing ratios are changing

nonlinearly. Finally, Prather (1997) pointed out that the

rate at which halogens are removed from the stratosphere

has its own time scale in addition to the lag time associated with transport of halocarbons into the stratosphere.

For short-lived halocarbons such as methyl bromide,

whose tropospheric abundance could decrease rapidly relative to the rates at which air is transported into or out of

the stratosphere, the decay of stratospheric inorganic

halogen resulting from a rapid drop in tropospheric burden

would be delayed and would not be well approximated by

a single lag time. This effect is less important for most

chlorinated ODSs because they have atmospheric lifetimes that are long relative to these transport rates.

Trends for EECl and EESC depend upon our understanding of the equivalency factor for bromine (alpha (a)),

which varies over different latitudes and altitudes and is

weighted by the distribution of stratospheric ozone

(Daniel et al., 1999; see also Section 1.4.4). Although 45

represents a globally-weighted mean for alpha, higher

estimates for this equivalency factor would suggest

smaller decreases for EECl during years in which bromine

continued to increase.

1.2.6 Fluorine in the Stratosphere

Measurements of hydrogen fluoride (HF) and carbonyl fluoride (COF2), which are primarily concentrated

in the stratosphere (Zander et al., 1992), provide an estimate of changes in the total fluorine burden of the

atmosphere (Sen et al., 1996). Although fluorine does not

catalytically destroy stratospheric ozone, trends in F*

y

(defined here as F*

y = [HF] + 2 ¥ [COF2]) have provided

an independent measure of changes in the abundance

of ozone-depleting substances (Gunson et al., 1994;

Anderson et al., 2000).

Since the previous Assessment (WMO, 1999), longterm investigations of HF and COF2 have continued, both

from the ground and from space. Related Jungfraujoch

results show a steady increase in column F*

y until the mid1990s, and a slower rate of increase afterward (Figure 1-

10; updates to Mahieu et al., 2000 and Mélen et al., 1998).

Two-dimensional (2-D) model calculations that include

atmospheric histories for the most abundant, anthropogenic F-containing gases (CFC-12, CFC-11, CFC-113,

HCFC-22, Halon-1211, and Halon-1301, which all are

important ODSs) reproduce the measured trends in F*

y

above the Junfraujoch station reasonably well (Figure 1-

10; Chipperfield et al., 1997). This consistency provides

further evidence that anthropogenic ODSs undergo photochemical degradation in the stratosphere and are responsible for the large increases in stratospheric chlorine and

fluorine observed during the past 20 years or more. Similar

conclusions have been drawn from analyses of the nearglobal mean HF measurements at 55 km from HALOE

(Anderson et al., 2000) and from ATMOS measurements

of individual F-containing gases between 1985 and 1994

over the altitude range of 8 to 60 km (Irion et al., 2002).

1.3 HALOCARBON SOURCES ESTIMATED

FROM INDUSTRIAL PRODUCTION 

The combination of halocarbon production data

from both the Alternative Fluorocarbons Environmental

Acceptability Study (AFEAS, 2001) and the United

Nations Environment Programme (UNEP, 2002) have

allowed for the first time a rigorous estimate of total global

production and consumption of ozone-depleting halocarSOURCE GASES

1.19
bons. Although the UNEP data provide more complete

geographical coverage than AFEAS, they are not audited.

The audited AFEAS data provide discrimination of production data by individual gases for all Parties other than

China, India, and Korea. For this report, the aggregated

additional production in these countries was subdivided

into individual compounds as follows: HCFCs were

treated as being produced in the same ratio as in the

AFEAS data, and only CFC-11 and -12 contributed to the

additional production of CFCs, consistent with there being

little production of CFC-113 outside of the companies

reporting to AFEAS (personal communication, M.R.

Harris, World Bank, 1999). In the case of methyl chloroform, global production data were derived from a similar

combination of an industrial database (McCulloch and

Midgley, 2001) and data in UNEP (2002). For the halons,

global production data were derived from Fraser et al.

(1999), updates to UNEP (1999), UNEP (2002), and also

from reports to the Parties on the Multilateral Fund, for

example, as described in Bingfeng et al. (2000).

None of these substances is emitted immediately

after it is produced; emissions occur over a period during

and after use that varies depending upon the application.

In some cases, notably where the substance is used as a solSOURCE GASES

1.20

Total C

olu

m

n

A

b

u

n

d

a

n

c

e (×

1

0

m

ole

c./c

m ) 15 2

1983.0 1986.0 1989.0 1992.0 1995.0 1998.0 2001.0 2004.0

2.5

2.0

1.5

1.0

0.5

0.0

COF2

HF

2D-Model F*

y

Inorganic fluorine above

Jungfraujoch

F*

y

June to November monthly means

Nonparametric LS fit (20%)

Figure 1-10. Time series of monthly mean vertical column abundances of HF (open circles: update from

Mahieu et al., 2000); COF2 (open triangles: update from Mélen et al., 1998); and their weighted sum, F*

y (filled

triangles; where F*

y = [HF] + 2×[COF2]) derived from observations at the Jungfraujoch (46.5°N) between 1983

and 2001. Only measurements from June to November are shown to avoid the significant variability observed

during winter and spring. The thick lines through the individual datasets represent nonparametric leastsquares fits with 20% weighting. Also shown is column F*

y above the Jungfraujoch calculated with a twodimensional model (Chipperfield et al., 1997). The model-calculated F*

y minima occur in late summer/fall. The

10-12% high bias in the model calculations compared with the summer/fall observations is believed to arise in

part from systematic uncertainties in the meridional transport of the 2-D model (see discussion in Chapter 4).

Long-term column HF measurements above Kitt Peak indicate rates of change similar to those observed at

the Jungfraujoch, but are not shown here (Rinsland et al., 2002b).
vent, emissions are prompt and can be well characterized

in quantity and timing. Since the previous Assessment

(WMO, 1999), emission functions that relate the quantities

of fluorocarbons in use to the quantities released into the

atmosphere have been intensively re-examined and, in

some cases, revised (Ashford, 2000). In the case of HCFC142b this has resulted in a significant change in the emission function and the calculated quantities emitted

(AFEAS, 2001). It is also apparent that containment of

most CFCs and HCFCs began to improve in the early

1990s. For example, emissions of CFC-11 from foam

blowing showed a sharp reduction after about 1992

(McCulloch et al., 2001), and these changes are included

in the revised release functions. Emissions from some

refrigeration applications have been reduced in recent years

(UNEP, 1998a; Calm, 1999), but these systems represent

only a small fraction of total emissions from refrigeration.

A single emission function for refrigeration has been used

for more than 30 years (McCulloch et al., 2001, 2002).

Ongoing work by the Technology and Economic

Assessment Panel (TEAP) has also allowed for independent estimates of halons, as discussed below.

1.3.1 CFCs and HCFCs

For substances used in refrigeration, there is an initial loss of between 5% (CFC-12) and 35% (HCFC-22),

and complete emission within 10 years for the average

system (Gamlen et al., 1986; Midgley and Fisher, 1993;

McCulloch et al., 2002). This is consistent with the

average linear loss rate of 11% yr-1 described by Baker

(1999). However, the actual distribution of emissions is

approximately normal around 4.5 ± 2 years.

1.3.2 Halons

Halons are primarily used as fire extinguishing

agents, although Halon-1301 also received minor use as

a cold-temperature refrigerant. UNEP (1998b) suggests

that production data as reported to the UNEP Ozone

Secretariat is accurate. Two estimates of emissions for

Halon-1211 and -1301 have been made using similar production data (Fraser et al., 1999; updated UNEP, 1999),

but since very different emission functions were used,

markedly different emission rates were obtained (see

Section 1.6). Further, the extent of inadvertent halon

emissions from some production facilities has yet to be

characterized but may be substantial (UNEP, 1999).

1.3.3 Carbon Tetrachloride (CCl4)

The principal identified use of carbon tetrachloride

is as feedstock for the production of CFC-11 and -12, and

from this process releases to the atmosphere arise only

from fugitive losses. Large discrepancies existed in

UNEP consumption data for past years owing to confusion over reporting procedures, and although these have

been addressed by TEAP (UNEP, 1998b), there may still

be significant omissions and unrecorded sources (see

Section 1.6). TEAP suggests emissions of 41(+50%,

-25%) Gg from total production of 203 Gg in 1996

(UNEP, 1998b), but, as discussed in Section 1.6, this is

only about half of the emissions necessary to account for

observed burdens and trends in that year given the updated

lifetime estimate of 26 years (see Section 1.4).

1.3.4 Methyl Chloroform (CH3CCl3)

Industrial production and release to the atmosphere

of methyl chloroform declined dramatically in the 1990s

(McCulloch and Midgley, 2001; UNEP, 2002). Global

anthropogenic emissions in 1999-2000 were estimated at

about 20 Gg yr-1, or only 3% of the emissions estimated

for 1989-1990, the peak emission years. Declines in emission closely follow declines in production because methyl

chloroform was used predominantly in applications where

releases occurred soon after sales.

Firn results (Butler et al., 1999; Sturrock et al.,

2002) suggest mixing ratios in the early 1900s, before

industrial production began, of less than 4 ppt and confirm that use by industry was the main source of atmospheric methyl chloroform.

Reports have suggested that small amounts of

methyl chloroform are emitted from biomass burning.

Upper limits quoted in an early study (Rudolph et al.,

1995) have since been revised downward with additional

laboratory work and suggest emissions may be on the

order of 2-10 Gg yr-1 from burning (Rudolph et al., 2000).

Emissions of this magnitude would sustain global mixing

ratios of 0.5-2 ppt, consistent with inferences from the

analysis of firn air.

1.4 HALOCARBON LIFETIMES, OZONE

DEPLETION POTENTIALS, AND GLOBAL

WARMING POTENTIALS 

1.4.1 Introduction

Halocarbon sinks chemically transform or physically remove halogen-containing source species from the

atmosphere after release. Known atmospheric sinks for

halocarbon source gases include photolysis, reaction with

hydroxyl radical (OH) in both the troposphere and stratosphere, reaction with electronically excited atomic oxygen

(O(1

D)) and atomic chlorine (Cl) chiefly in the stratoSOURCE GASES

1.21
sphere, uptake in oceanic surface waters through chemical and biological degradation processes, biological

degradation in soils, and possibly surface reactions on

minerals. Not all halocarbon source gases become

degraded by all of these sink processes. The quantitative

determinations of trace gas sink strengths are of interest

because the sinks control both the atmospheric lifetimes

and the location of halogen release from source gases.

The important sink processes for atmospheric lifetime are those that have significant strength on a mass or

molecule basis, that is, where loss frequency and atmospheric abundance are co-located. A simple definition of

lifetime is the sum of the integrals over the entire atmosphere of the product of the first-order sink process frequencies (the local instantaneous lifetimes) and the

normalized source gas mass or concentration distribution.

1.4.2 Lifetimes

For a given halocarbon, each relevant sink process

contributes to the additive first-order total loss frequency,

L, which is variable in space and time. A local halocarbon

lifetime, tlocal, can be defined as the inverse of L evaluated at a point in space and time:

(1-1)

For consideration of global effects of the halocarbon, L

must be integrated globally, and the integral must be

weighted by the distribution of the halocarbon on which

the sink processes act. If we consider a normalized halocarbon distribution n(x, y, z, t), a global instantaneous lifetime can be defined as

(1-2)

where V is an atmospheric volume element. This expression can be averaged over a year to determine the global,

annually averaged lifetime. Because the total loss frequency L is the sum of the individual sink process frequencies, tglobal can also be expressed in terms of process

lifetimes; for example,

Note that in considering lifetime with respect to

individual sink processes, which may be defined or analyzed over limited domains, the associated burden (the

numerator in Equation (1-2)), however, must always be

global and include all communicating reservoirs. When

the halocarbon distribution is time invariant, where emissions, sinks, and atmospheric transport are in long-term

balance, the halocarbon burden B(t) can be defined by

(1-4)

and the lifetime tglobal becomes a scale factor relating the

halocarbon emission rate E(t) (or equivalently the sink

rate) to the atmospheric burden. This steady-state lifetime is also formally equivalent to the scale factor relating

the time-integrated atmospheric burden to a release of

halocarbon at a single point in time (Prather, 1996). For

source compounds that do not persist long enough to

become well mixed in the troposphere (here assumed for

gases with t < 0.5 years), distributions are strongly

dependent on the pattern of emissions. Consequently, the

concept of a single global lifetime, ODP, or GWP becomes

inappropriate for such short-lived gases (see Chapter 2).

In this chapter, only local lifetimes are reported for shortlived gases (Table 1-3), and these lifetimes are not derived

from any actual steady-state global emission/distribution

patterns. (GWPs for these gases are not reported in Table

1-6 because of these concerns.) Finally, a discussion of

the ODP concept for short-lived gases is addressed in

Chapter 2.

WMO (1999) and IPCC (2001) include detailed

discussions of the various specific techniques for estimating steady-state and non-steady-state global lifetimes,

to which the reader is referred. The techniques fall into

four general categories: (1) lifetimes derived from modelcalculated distributions and sink loss frequencies, (2) lifetimes inferred by inverse methods from direct observations of halocarbon abundance with high precision and

accuracy over multiyear periods and requiring varying

degrees of knowledge about actual emissions, (3) lifetimes determined relative to a reference species with an

independently derived lifetime through observations of

species correlations or knowledge of relative kinetic rate

parameters, and (4) lifetimes determined from some combination of the above three techniques, where the total

lifetime is determined as the combination of individual

process lifetimes.

Important advancements have been made in our

understanding of halocarbon lifetimes since WMO (1999).

Those relate to improvements in understanding the lifetime of methyl chloroform and in choosing a globally

representative temperature that allows better estimates of

lifetimes of other gases destroyed by OH.

Recently reported analyses of methyl chloroform

lifetime based on observations from global sampling netSOURCE GASES

1.22

t local = 1

Lxyzt (, ,,)

t global = Ú =

Ú

dV n V t

dV L V t n V t

( , )( )

(, ) (, )

1

t tt t t global

OH ocean photolysis other processes

=+ + +

È

Î

 ̆

 ̊

-

11 1 1 dB t

dt E t ( ) B t ( ) ( ) = - t global

(1-3)
works over multiyear periods (Montzka et al., 2000; Prinn

et al., 2000, 2001) suggest a slightly longer central value

than previous World Meteorological Organization (WMO)

and Intergovernmental Panel on Climate Change (IPCC)

reports (Kurylo and Rodríguez et al., 1999; Prather and

Ehhalt et al., 2001). Additionally, stratospheric in situ correlation measurements (Volk et al., 1997) derive a somewhat shorter stratospheric lifetime than was adopted in

these Assessments. Because CH3CCl3 acts as the reference compound for the relative determination of lifetime

with respect to tropospheric OH oxidation for the large

suite of hydrogenated halocarbons, changes in CH3CCl3

lifetime need to be carefully considered. Equation (1-3) is

employed to determine the value of tOH for CH3CCl3 (here

derived to be 6.1 years, see below), where tglobal ( = 5.0

years) is inferred from direct observations of CH3CCl3

and estimates of emissions; tphotolysis ( = 38-41 years) arises

from loss in the stratosphere and is inferred from observed

stratospheric correlations among CH3CCl3, CFC-11, and

the stratospheric age of the air mass observed; tocean

( = 94 years) is derived from a model of the oceanic loss

process; and tother processes is taken to be zero.

The total atmospheric lifetime value adopted here

of 5.0 years for methyl chloroform at steady state (i.e.,

unchanging distribution or balanced emission and loss)

lies between the mean lifetimes inferred from observations for non-steady-state periods of 5.2 (+0.2/-0.3) years

(Montzka et al., 2000) and 4.9 (+0.6/-0.5) years (Prinn et

al., 2001) and below the observationally derived upper

limit of 5.5 years (Montzka et al., 2000). Note that the

difference between these observational lifetimes is not

ascribed to simple uncertainties in calibration accuracy or

measurement precision. The Prinn et al. (2001) value pertains to the average global OH during 1978-2000, whereas

the Montzka et al. (2000) value refers specifically to 1998-

1999, a time period for which Prinn et al. (2001) suggest

that global OH fell below the long-term average, producing a longer instantaneous lifetime.

The oceanic sink lifetime of methyl chloroform has

been reassessed recently based on a high-resolution model

of ocean processes and surface mixing, and a value of 94

years was calculated (Yvon-Lewis and Butler, 2002), similar to the lifetime adopted in the previous Assessment

(WMO, 1999). Given the above-mentioned revisions, the

previously adopted process OH lifetime of 5.7 years would

require an unrealistically long stratospheric-loss lifetime

(much greater than 45 years), where recent reports indicate even shorter values (≤40 years). A longer lifetime of

6.1 years with respect to tropospheric hydroxyl radical

reaction is consistent with stratospheric-loss lifetimes of

38 to 41 years. This range of stratospheric-loss lifetime

lies between lower values inferred from tracer-tracer correlations (Volk et al., 1997; Schauffler et al., 2002) and

higher values calculated in multidimensional models

(WMO, 1999). The general effect of this updated recommendation is to lengthen lifetimes of hydrogen-containing

halocarbons and methane by about 7% relative to the two

most recent Assessment reports, and to increase estimates

for their ODP and GWP values by a similar amount.

Spivakovsky et al. (2000) have computed a global

climatological distribution of tropospheric OH based on

observed distributions of the OH-controlling quantities.

Using this OH distribution and assuming a well-mixed

distribution for the OH reaction partner, they show that

selecting a temperature of 272 K to evaluate a tropospheric

average OH rate constant produces a good fit over a wide

range of activation energies in the Arrhenius rate constant

expression. As in IPCC (2001), this value is adopted here

in calculating lifetimes by reference to CH3CCl3 in Table

1-3. WMO (1999) used an earlier 277 K estimate for this

value. The OH reaction rate constants used in generating

the Table 1-3 lifetimes are the new recommendations of

the NASA Panel for Data Evaluation (Sander et al., 2002).

The global lifetime for Halon-1211 (CBrClF2) has

been reassessed. This gas is destroyed exclusively by photolysis in both the troposphere and stratosphere, which

makes it a somewhat special case among halogen source

gases. Past lifetime recommendations that depend on

cross section data and model calculations include 12-18

years (WMO, 1989), 18.5-20.1 years (WMO, 1991), 9.3-

28.2 years (Kaye et al., 1994), and 20 years (WMO, 1995).

The Halon-1211 lifetime has also been inferred from a

comparison between observed atmospheric burdens and

emission histories derived from halon production magnitudes. Uncertainty in both the atmospheric measurements

and available release histories (Fraser et al., 1999; an

update to UNEP, 1999) do not constrain these estimates

to better than 11-30 years (see Section 1.6). Finally, the

stratospheric component of the Halon-1211 lifetime has

been inferred from correlating stratospheric observations

of this halon to age of stratospheric air, nitrous oxide

(N2O), and CFC-11 (e.g., Volk et al., 1997). This stratospheric lifetime is difficult to combine appropriately with

modeled estimates of tropospheric loss, although the

available estimates suggest a global Halon-1211 lifetime

of about 14 years. Because of the uncertainties surrounding observational-based lifetime estimates, the recommendation for the total steady-state lifetime of Halon1211 is updated here to 16 years based primarily on

reported model-calculated values known to include the

tail of the Halon-1211 absorption between 300 and 400

nm (Burkholder et al., 1991; Fraser et al., 1999). Clearly

additional laboratory, observational, and modeling studies

will be needed to reduce the uncertainties in this lifetime.

SOURCE GASES

1.23
SOURCE GASES

1.24

Table 1-3. Halogenated trace gas lifetimes.

Industrial Designation Chemical Formula Lifetime (years) Notes

or Common Name

Halogen-substituted methanes

HFC-41 CH3F 2.4 1

HFC-32 CH2F2 4.9 1

HFC-23 CHF3 270 1

FC-14 (carbon tetrafluoride) CF4 50 000 2

Methyl chloride CH3Cl 1.3 1, 3

Dichloromethane CH2Cl2 0.38 15, 16

Chloroform CHCl3 0.41 15, 16

Carbon tetrachloride CCl4 26 5

HCFC-31 CH2ClF 1.3 4

HCFC-22 CHClF2 12.0 1, 6

HCFC-21 CHCl2F 1.7 1, 6

CFC-13 CClF3 640 2

CFC-12 CCl2F2 100 2, 7

CFC-11 CCl3F 45 2, 7

Methyl bromide CH3Br 0.7 3

Dibromomethane CH2Br2 0.33 8, 15, 16

Bromoform CHBr3 0.07 8, 15, 16

Bromodifluoromethane CHBrF2 5.8 4

Bromochloromethane CH2BrCl 0.37 8, 15, 16

Bromodichloromethane CHBrCl2 0.21 7, 15, 1

Dibromochloromethane CHBr2Cl 0.19 7, 15, 16

Halon-1301 CBrF3 65 2, 7

Halon-1211 CBrClF2 16 9

Halon-1202 CBr2F2 2.9 17

Methyl iodide CH3I 0.02 7, 15, 16

Diiodomethane CH2I2 Minutes 7, 15, 16

Chloroiodomethane CH2ClI Hours 7, 15, 16

Trifluoroiodomethane CF3I <0.005 2, 15

Halogen-substituted ethanes

HFC-161 CH3CH2F 0.21 2, 15

HFC-152 CH2FCH2F 0.60 2, 15

HFC-152a CH3CHF2 1.4 1, 6

HFC-143 CH2FCHF2 3.5 1

HFC-143a CH3CF3 52 1

HFC-134 CHF2CHF2 9.6 1

HFC-134a CH2FCF3 14.0 1, 6

HFC-125 CHF2CF3 29 1, 6

FC-116 (perfluoroethane) CF3CF3 10 000 2

chloroethane CH3CH2Cl 0.11 15

1,1 dichloroethane CH2ClCH2Cl 0.19 10, 15

Methyl chloroform CH3CCl3 5.0 9
SOURCE GASES

1.25

Table 1-3, continued.

Industrial Designation Chemical Formula Lifetime (years) Notes

or Common Name

HCFC-142b CH3CClF2 17.9 1, 6

HCFC-141b CH3CCl2F 9.3 1, 6

HCFC-123 CHCl2CF3 1.3 6, 11

HCFC-124 CHClFCF3 5.8 6, 11

CFC-113 CCl2FCClF2 85 2

CFC-113a CCl3CF3 NA NA

CFC-114 CClF2CClF2 300 2

CFC-115 CClF2CF3 1700 2

Halon-2402 CBrF2CBrF2 20 17

Halogen-substituted propanes

HFC-281ea CH3CHFCH3 0.06 15

HFC-263fb CH3CH2CF3 1.6 2

HFC-245ca CH2FCF2CHF2 6.2 1

HFC-245ea CHF2CHFCHF2 4.0 2

HFC-245eb CH2FCHFCF3 4.0 12

HFC-245fa CHF2CH2CF3 7.6 1

HFC-236cb CH2FCF2CF3 13.6 1

HFC-236ea CHF2CHFCF3 10.7 1

HFC-236fa CF3CH2CF3 240 1

HFC-227ea CF3CHFCF3 34.2 1

FC-218 (perfluoropropane) CF3CF2CF3 2600 2

n-propyl chloride CH3CH2CH2Cl 0.06 13, 15

HCFC-243cc CH3CF2CFCl2 26.4 4

HCFC-225ca CHCl2CF2CF3 1.9 11

HCFC-225cb CHClFCF2CClF2 5.8 11

n-propyl bromide CH3CH2CH2Br 0.04 15, 16

n-propyl iodide CH3CH2CH2I 0.003 7, 15, 16

isopropyl iodide CH3CHICH3 0.002 7, 15

Halogen-substituted higher alkanes

HFC-365mfc CH3CF2CH2CF3 8.6 1

HFC-356mcf CH2FCH2CF2CF3 1.2 1

HFC-356mff CF3CH2CH2CF3 8.1 1

HFC-338pcc CHF2CF2CF2CHF2 12.3 1

FC-318 (perfluorocyclobutane) c-C4F8 3200 1

FC-31-10 (perfluorobutane) C4F10 2600 2

HFC-43-10mee CF3CHFCHFCF2CF3 15.9 1

HFC-458mfcf CF3CH2CF2CH2CF3 23.2 1

FC-41-12 (perfluoropentane) C5F12 4100 2

HFC-55-10mcff CF3CF2CH2CH2CF2CF3 7.7 2

FC-51-14 (perfluorohexane) C6F14 3200 2
SOURCE GASES

1.26

Table 1-3, continued.

Industrial Designation Chemical Formula Lifetime (years) Notes

or Common Name

Fluorinated alcohols

CF3CH2OH 0.41 15

CF3CF2CH2OH 0.39 15

(CF3)2CHOH 2.0 4

Fluorinated ethers

HFE-152a CH3OCHF2 1.6 1

HFE-143a CH3OCF3 4.3 1

HFE-134 CHF2OCHF2 26 1

HFE-125 CHF2OCF3 136 1

HFE-227ea CF3OCHFCF3 11 2

HFCE-235da2 CHF2OCHClCF3 2.6 2

HFE-236ea2 CHF2OCHFCF3 5.8 2

HFE-236fa CF3OCH2CF3 3.7 2

HFE-245fa1 CHF2CH2OCF3 2.2 2

HFE-245fa2 CHF2OCH2CF3 4.9 1

HFE-245cb2 CH3OCF2CF3 5.1 4

HFE-254cb2 CH3OCF2CHF2 2.6 4

HFE-263fb2 CH3OCH2CF3 0.1 2

HFE-329mcc2 CF3CF2OCF2CHF2 6.8 2

HFE-338mcf2 CF3CF2OCH2CF3 4.3 2

HFE-347mcc3 CH3OCF2CF2CF3 5.2 4

HFE-347mcf2 CF3CF2OCH2CHF2 2.8 2

HFE-347pf3 CHF2OCH2CF2CF3 5.9 4

HFE-347sy2 CF3CF(OCH3)CF3 3.7 4

HFE-356mec3 CH3OCF2CHFCF3 0.94 2

HFE-356mff2 CF3CH2OCH2CF3 0.4 15

HFE-356pcc3 CH3OCF2CF2CHF2 0.93 2

HFE-356pcf3 CHF2OCH2CF2CHF2 3.6 4

HFE-356pcf2 CHF2CH2OCF2CHF2 2 2

HFE-365mcf3 CH3OCH2CF2CF3 0.11 2, 15

HFE-374pc2 a CH3CH2OCF2CHF2 5 2

CF3CH(OCF3)CHF2 3.1 2

(CF3)2CFOCH3 3.4 2

HFE-7100 C4F9OCH3 5 2

HFE-7200 C4F9OC2H5 0.77 2

H-Galden 1040x b CHF2OCF2OCF2CF2OCHF2 6.3 2

HFE-236ca12 CHF2OCF2OCHF2 12.1 2

HFE-338pcc13 CHF2OCF2CF2OCHF2 6.2 2

Other fluorinated species

Trifluoromethylsulfurpentafluoride SF5CF3 800 18

Sulfur hexafluoride SF6 3200 2
Previous lifetime estimates of carbon tetrachloride

were derived from observations and modeling of photolytic loss rates in the stratosphere (Volk et al., 1997;

Prinn and Zander et al., 1999). Reports of widespread

observations of undersaturation of carbon tetrachloride in

the ocean (Huhn et al., 2001; Wallace et al., 1994) suggest an additional, significant loss for this gas (YvonLewis and Butler, 2002). When combined with the other

loss processes, a revised global lifetime of 26 (17-36)

years is now calculated for this gas (Yvon-Lewis and

Butler, 2002).

1.4.3 Fractional Release Factors

In considering the effect of halogen source gases

on ozone, it is first necessary to determine the composition of tropospheric air entering the stratosphere, mostly

through the tropical tropopause. For compounds with lifetimes measured in years, the amount of halogen entering

the stratosphere for a specific level of emission is

inversely related to the lifetime at steady state. Once a

halogen source gas is in the stratosphere, release of a

halogen atom from the source gas through photolysis or

chemical reaction is a function of stratospheric local lifetime or loss frequency and differs greatly from compound

to compound. More complete release of the halogen atom

catalyst in the stratosphere produces a greater extent of

local photochemical loss of ozone, for a given source gas

or in comparing source gases, all else being equal.

A point measurement of a given halogen source

molecule in the stratosphere can be recast as a fractional

release (FR), defined as

(1-5)

where r is the source compound mixing ratio or mole fracSOURCE GASES

1.27

Table 1-3, continued.

Footnotes

For completeness, estimates for local lifetimes for some very short-lived (t < 0.5 years) species are included. As discussed in Chapter 2, the atmospheric lifetimes for these species (defined as the ratio of burden to emission; see Prather and Ehhalt et al., 2001) depend on the location and time of

emission. Thus, these local lifetimes should not be used in estimation of semiempirical ODP, GWP, or EESC calculations for these gases.

Notes:

1. Global lifetime estimated from a process lifetime with respect to tropospheric OH calculated relative to 6.1 years for CH3CCl3, assuming an

average temperature of 272 K (Spivakovsky et al., 2000; Prather and Ehhalt et al., 2001); OH rate constants from Sander et al. (2002); and stratospheric loss lifetimes inferred from IPCC (2001).

2. Prather and Ehhalt et al. (2001) and Ramaswamy et al. (2001).

3. See Section 1.5 for further discussion related to methyl halide global lifetimes. 

4. Lifetime calculated as in Note 1 except that no estimate of a stratospheric loss lifetime was available to include in the lifetime estimate listed.

Hence this is an upper bound to the global lifetime estimate.

5. IPCC (2001) and including an oceanic loss term with 94-year lifetime observed in saturation data and ascribed to an unidentified process (YvonLewis and Butler, 2002).

6. Including oceanic loss term from Yvon-Lewis and Butler (2002). The contribution of oceanic loss to the lifetime of HCFC-21, HCFC-22, HCFC123, HCFC-124, HCFC-141b, HFC-125, and HFC-152a is small; for HFC-134a and HCFC-142b it is negligibly small at the reported precision.

7. WMO (1999).

8. Lifetimes listed include local tropospheric photolysis lifetimes from Table 2-9 in Kurylo and Rodríguez et al. (1999). Consideration of only tropospheric OH loss results in local lifetimes of 0.34 years for CH2Br2, 0.21 years for CHBr3, and 0.37 years for CH2BrCl. 

9. See Section 1.4 text for discussion.

10. OH rate constant from Qui et al. (1992).

11. Lifetime calculated as in Note 1, but with stratospheric loss from Naik et al. (2000).

12. Lifetime calculated as in Note 1, but with OH rate constant and stratospheric loss from Naik et al. (2000).

13. Markert and Nielsen (1992).

14. OH rate constant from DeMore and Bayes (1999).

15. The values estimated correspond to local lifetimes in the free troposphere. For species that react with OH, the process lifetime due to OH reaction

is calculated using the rate constant at 275 K (for lifetimes greater than 10 days) or 300 K (for lifetimes less than 10 days) and OH concentration of

1 ¥ 106 molec cm–3. These should not be used in estimating ODP, GWP, or EESC because the atmospheric burden for these short-lived gases (t <

0.5 years) depends on the location and time of emissions.

16. See Chapter 2.

17. From the 2-D model calculation in Fraser et al. (1999).

18. Takahashi et al. (2002).

a Referred to as HFE-374pcf2 in past Assessments.

b Also known as HFE-43-10pccc124.

FR entry point

entry

=

r r -

r
tion. Observations of halogen source compounds at high

spatial resolution in the stratosphere, chiefly by in situ

techniques, show a compact relationship when plotted as

a scatter correlation of fractional release of compound x

to that of compound y. Identifying compound y as CFC11 produces the fractional release factor that is used in the

semiempirical determination of ODPs. It is essentially a

measure of the local release or availability of the inorganic halogen compounds that actually produce the ozone

depletion, relative point by point to CFC-11. These relative fractional release factors are used in the calculation

of ODP, EECl, and EESC.

Although the correlations are compact, they are

typically not simply linear, because local loss frequencies

differ from compound to compound. Conceptually, the

fractional release factor could be globally integrated,

weighted by the local ozone loss. In practice, Schauffler

et al. (1999) suggest the use of a restricted range of the

correlation representing mid-high latitudes, where the

greatest loss of ozone occurs. Schauffler et al. (2002)

report fractional release factors for three ranges of stratospheric age of air, G, determined from N2O abundance

observed during the Stratospheric Aerosol and Gas

Experiment (SAGE) III Ozone Loss and Validation

Experiment (SOLVE), Stratospheric Tracers of

Atmospheric Transport (STRAT), and Photochemistry of

Ozone Loss in the Arctic Region in Summer (POLARIS)

aircraft campaigns. High values of N2O, near those in the

troposphere, are associated with air that has recently

entered the stratosphere, whereas low N2O values are

indicative of air that has cycled through the stratosphere

by the mean circulation. The mean halocarbon fractional

release values associated with ages of 2.0-4.5 years are

appropriate for the midlatitude middle stratosphere.

Earlier studies (Daniel et al., 1995) suggest a reference

CFC-11 fractional release value of 0.8 for air that had

passed through the mean circulation. Schauffler et al.

(2002) relate this to older stratospheric air (G = 4.5 years).

For 3.5-yr-old air, they indicate a CFC-11 fractional

release of 0.55. Relative fractional release values derived

from stratospheric observations of air of intermediate ages

are given in Table 1-4. Because these are relative quantities, they are sensitive to nonlinearities in the slopes of

correlation plots of release against age of air, but less sensitive to the absolute CFC-11 reference value.

Models and a semiempirical framework have been

used to estimate fractional release factors for unmeasured

or currently nonexistent atmospheric source gases

(Solomon et al., 1992). Models predict ozone loss explicitly, however, so it is generally more straightforward to

integrate the modeled ozone loss and calculated ODP

directly as opposed to using a modeled release factor in

the semiempirical approach. The fractional release values

for HCFC-141b and -142b used here and in past

Assessments (Table 1-4) are based upon a semiempirical

analysis of measured stratospheric profiles for methane

and other gases (Solomon et al., 1992). A recent analysis

of HCFC observations in the stratosphere, however, suggests the possibility for much smaller (by factors of 3 to

5) fractional release values and ODPs for HCFC-141b and

-142b (Schauffler et al., 2002; Tables 1-4 and 1-5). Such

low estimates are difficult to reconcile with the earlier

estimates, given that only comparatively small refinements have been noted for reaction rate constants and

other relevant model parameters since the original calculations were performed. Without additional research,

changes to ODPs calculated for these HCFCs cannot be

recommended at this time.

1.4.4 Ozone Depletion Potential

The Ozone Depletion Potential (ODP) is an integrative quantity, distinct for each halocarbon source

species, that represents the extent of ozone depletion in

the stratosphere expected from the halocarbon on a massfor-mass basis relative to CFC-11. The formal definition

of ODP is the ratio of integrated perturbations to total

ozone, for a differential mass emission of a particular

compound relative to an equal emission of CFC-11. An

extensive discussion of the details of the ODP calculation

is given in WMO (1995).

Solomon and coworkers (Solomon et al., 1992;

WMO, 1995) introduced a semiempirical ODP formulation that takes advantage of the relative nature of the ODP

definition to remove the requirement for explicit representations of ozone-controlling photochemistry in the

context of chemistry-transport models. The quantities

required are physical properties of the halocarbons, their

lifetimes, a factor representing the distribution of inorganic halogen release in the stratosphere obtained from

observation, and, in the case of bromine and iodine, an

analysis of the catalytic efficiency of ozone destruction

relative to chlorine, derived from knowledge of the kinetic

parameters of the partitioning of the halogen radical

families. For long-lived gases that are well mixed in the

troposphere, the semiempirical ODP definition is

where the fractional release factors are discussed in Section

1.4.3 above, a is the relative effectiveness of any halogen

compared with chlorine for ozone destruction (taken to be

SOURCE GASES

1.28

OPD

fractional release factor

CFC-11

CFC 11

( )

( )

x M

M

x n

x

x

=

◊◊ ◊ ◊ - a t

t 3 (1-6)
45 here for bromine; Daniel et al., 1999), tx is the global

lifetime of the long-lived, well-mixed gas, Mx is the molecular weight, and nx is the number of halogen atoms.

Daniel et al. (1999) have discussed, specifically for

bromine, the definitions and determination of the effectiveness parameter, a, of a halogen radical family in catalyzing ozone destruction, relative to the chlorine radical

family. Their suggested definition of a, to be most useful

in the ODP definition given in Equation (1-6) above, is

cast in terms of the column response of ozone perturbation

to inorganic halogen change. Although appearing as a

single, fixed, global quantity, it is actually an integrated

result with sensitivity to both kinetic parameters for chlorine species and the other halogen of interest, and to atmospheric transport behavior. For this reason, there is no

simple way to adjust a quantitatively to reflect new determinations, for example, of individual rate coefficients.

Because a single global value is required by the

ODP definition, a represents the global correlation of

local photochemical behaviors, for the catalyst of interest

(e.g., bromine) and chlorine, with the distribution of ozone

loss. Relative effectiveness for localized regions may

differ substantially. For example, Chipperfield and Pyle

(1998) model a value of 60 for bromine in the Arctic

vortex. The value of 45 used here represents a midlatitude result, which is intended to be more representative of

the global mean relative to ozone depletion. An important point made by Danlin et al. (1996) and Chipperfield

and Pyle (1998) is that the relative effectiveness, a, can

be sensitive to the atmospheric state. Bromine is more

effective relative to chlorine in the Arctic vortex for higher

inorganic chlorine levels in the range of 2-4 ppb.

Semiempirical ODP values calculated for the current

atmosphere will be different from values calculated for a

future atmosphere containing less chlorine and altered

distributions of ozone loss.

Daniel et al. (1999) based their result on DeMore

et al. (1997), which did not make a recommendation for

products of the OH + chlorine monoxide (ClO) reaction,

but discussed studies showing the predominant product

channel is hydroperoxyl radical (HO2) + Cl. Sander et al.

(2000) do recommend a value for the product channel

SOURCE GASES

1.29

Table 1-4. Fractional release factors relative to CFC-11 (CCl3F).

Compound Relative Fractional Release Factor a Schauffler et al. (2002) b Notes

CFC-11 1 By definition

CFC-12 0.60 0.57 ± 0.05 1

CFC-113 0.75 0.65 ± 0.05 1

CFC-114 0.28 ± 0.02

CFC-114a 0.63 ± 0.10

CCl4 1.06 1.14 ± 0.06 1

HCFC-22 0.35 0.29 ± 0.02 1

CH3CCl3 1.08 1.29 ± 0.16 1

HCFC-123 1.11 1

HCFC-124 0.52 1

HCFC-141b 0.72 0.23 ± 0.11 1

HCFC-142b 0.36 0.08 ± 0.04 1

HCFC-225ca 1.1 2

HCFC-225cb 0.5 2

CH3Cl 0.80 0.85 ± 0.03 2, 3

CH3Br 1.12 4

Halon-1301 0.62 4

Halon-2402 1.22 4

Halon-1211 1.18 1.24 ± 0.07 4

Notes for Relative Fractional Release Factors in column 2:

1. Daniel et al. (1995).

2. WMO (1999).

3. Solomon et al. (1992).

4. Schauffler et al. (1999).

a These fractional release estimates are used after multiplication by 0.8 (to account for absolute fractional release of CFC-11) in the calculation of

EESC (Table 1-17; Section 1.8). See Notes above for references.

b Estimated for stratospheric air having a mean age of 2-4.5 years.
forming HCl, based on Lipson et al. (1999). Including

this product channel can substantially diminish the modeled ClO abundance in the mid-stratosphere, by as much

as a factor of 2. In reducing the fraction of inorganic chlorine present in the active ClO form, the new recommendation has the effect of moderating the chlorine-induced

ozone change in that region, and consequently increasing

the modeled efficiency of bromine relative to chlorine.

Quantifying the change would require a comprehensive

reanalysis, but because the region most affected contributes a relatively small portion of the column ozone

change, the overall value for a should not change greatly,

and the estimate of 45 from Daniel et al. (1999) remains

the best available at the present time.

To separate their study of the effectiveness parameter, a, from the effects of halogen fractional release distributions for any source gas, Daniel et al. (1999) assumed

a hypothetical bromine source species characterized by

the CFC-11 fractional release pattern. They point out that

sources with significantly shorter atmospheric lifetimes

than CFC-11 would also be characterized by greater fractional release lower in the stratosphere, which would produce a larger effectiveness parameter value. The effect

for iodine source species would be even more pronounced

than for bromine sources. This interaction between the

fractional release parameter and the effectiveness parameter introduces uncertainty into the ODP estimates and

should be analyzed further for new source species of

interest and in the light of projected changes in future stratospheric chlorine loading.

Table 1-5 includes both model and semiempirical

ODP values. Semiempirical values are calculated using

SOURCE GASES

1.30

Table 1-5. Updated Ozone Depletion Potentials for long-lived halocarbons.

Halocarbon Updated Updated Semiempirical WMO (1999) WMO (1999) Montreal

Model- Semiempirical b Schauffler et al. Model Semiempirical Protocol

Derived a (2002) c

CFC-11 1

CFC-12 1.0 0.96 0.82 0.9 1.0

CFC-113 1.0 0.90 0.90 0.9 0.8

CFC-114 0.94 1.00 0.85 1.0

CFC-115 0.44 0.40 0.6

Halon-1301 12 12 13 10.0

Halon-1211 6.0 6.0 5.1 5 3.0

Halon-2402 < 8.6 6.0

Halon-1202 1.3 d

CCl4 0.73 0.78 1.20 1.1

CH3CCl3 0.12 0.15 0.11 0.12 0.1

HCFC-22 0.05 0.041 0.034 0.05 0.055

HCFC-123 0.02 0.012 0.02 0.02

HCFC-124 0.02 0.026 0.022

HCFC-141b 0.12 0.037 0.086 0.1 0.11

HCFC-142b 0.07 0.014 0.043 0.066 0.065

HCFC-225ca 0.02 0.017 0.025 0.025

HCFC-225cb 0.03 0.017 0.03 0.033

CH3Cl 0.02

CH3Br 0.38 0.37 0.37 e 0.6

Upper limits for selected hydrofluorocarbons

HFC-134a < 1.5 ¥ 10–5

HFC-23 < 4 ¥ 10–4

HFC-125 < 3 ¥ 10–5

a Calculated from ODPs reported in Table 11-1 of WMO (1999) with lifetimes from Table 1-3 in this chapter.

b Calculated with updated lifetimes (Table 1-3) and fractional release factors (column 2, Table 1-4) except in the case of Halon-1202. c Calculated with updated lifetimes (Table 1-3) and fractional release factors from Schauffler et al. (2002) (column 3, Table 1-4).

d Derived from the semiempirical estimate in WMO (1991) by considering updated lifetimes for Halon-1202 and CFC-11 (Table 1-3) and an alpha of 45. e Value was erroneously reported in Table 11-1 of WMO (1999) as 0.57.
updated lifetimes and two separate estimates of fractional

release factors (listed in Table 1-4). The values for CFC12 and CFC-113 are updated from the semiempirical

ODPs in Table 11-1 of WMO (1999), which were based

on the reference lifetimes of WMO (1995), not WMO

(1999). The value of 0.38 for CH3Br is based on a total

lifetime of 0.7 years and is essentially unchanged from

the semiempirical value given in Table 2-6 of the previous

Assessment (Kurylo and Rodríguez et al., 1999) and the

model-derived value reported in Table 11-1 of that

Assessment (Madronich and Velders et al., 1999). (Note

that the semiempirical value given in Table 11-1 of WMO

(1999) was incorrectly recorded.) ODPs derived for

HCFC-141b and -142b from the new observation-based

fractional release factors are much lower than those

derived from earlier model calculations (Schauffler et al.,

2002; Solomon et al., 1992). Further work will be required

to resolve these discrepancies.

1.4.5 Global Warming Potential

Global Warming Potentials (GWPs) are used to

estimate the integrated climate forcing of various greenhouse gases compared with a reference gas, usually

chosen to be carbon dioxide (CO2). They provide a simple

way to gauge how decisions affecting greenhouse gas

emissions (e.g., in the Kyoto Protocol) may influence our

future climate in a relative sense (see, e.g., Ramaswamy

et al., 2001, and references therein). Furthermore they

allow the effects of these potential actions to be roughly

assessed in terms of a “carbon dioxide equivalent” emission. Radiative effects of CFC alternatives are also

assessed with the use of GWPs.

Because of the recent publication of Climate

Change 2001: The Scientific Basis (IPCC, 2001), we only

report updates to that work here. We refer the reader to

Chapter 6 of IPCC (2001) for a thorough discussion of

radiative forcing and GWPs.

The GWP of a particular compound is defined as

the ratio of the time-integrated radiative forcing from the

instantaneous emission of 1 kg of some gas relative to

that of 1 kg of a reference gas. Mathematically, the GWP

is given by

(1-7)

where ax is the radiative forcing of a unit mass of species

x added to the current atmospheric composition, x(t) is the

atmospheric decay function of the pulse of species x, ar

and r(t) are the corresponding quantities for the reference

gas, and TH is the time horizon over which the calculation is performed. The response function used in this

chapter is the same as in Granier and Shine et al. (1999)

and IPCC (2001). The formula for calculating r(t) where

CO2 is the reference gas is also the same as in Granier and

Shine et al. (1999) and Ramaswamy et al. (2001). We

have adopted a present-day mixing ratio of atmospheric

CO2 of 370 ppm (mmol mol-1; see Section 1.7.1), higher

than that of both Granier and Shine et al. (1999) and

Ramaswamy et al. (2001) (364 ppm), reflecting the continued increasing atmospheric abundance of this

radiatively important gas. This increased CO2 abundance,

relative to Ramaswamy et al. (2001), leads to a decrease

in the radiative forcing of a small addition of atmospheric

CO2, and a corresponding increase in all GWPs, of about

1.3%.

1.4.5.1 RADIATIVE FORCING UPDATES

Since Granier and Shine et al. (1999), three comprehensive modeling studies have been published that

estimate the radiative forcing efficiency of many radiatively important gases. Two of these, Highwood and

Shine (2000) and Jain et al. (2000), were considered in

Ramaswamy et al. (2001), and the third, Sihra et al.

(2001), was not published until after Ramaswamy et al.

(2001). For the majority of gases appearing in Tables 1-6

and 1-7, the radiative forcing efficiencies calculated by

Sihra et al. differ from the adopted Ramaswamy et al.

(2001) values by less than 20%; for all but one of these

gases, we retain the forcing efficiencies adopted for

Ramaswamy et al. (2001). HCFC-21 is changed because

of an error in its tabulated value in Granier and Shine et

al. (1999) and Ramaswamy et al. (2001) (see notes to

Table 1-6). Radiative efficiencies for six gases differed

by more than 20% between the Sihra et al. (2001) work

and Ramaswamy et al. (2001). Revisions to radiative

forcing efficiencies for two of these gases (HCFC-123

and HCFC-225ca) are recommended and are discussed in

the notes to Table 1-6. For the other four we retain the

Ramaswamy et al. (2001) recommendation either because

there was no other corroborative calculation to go along

with the Sihra et al. (2001) result (i.e., for HFE-143a and

H-Galden 1040x) or because the Sihra et al. values differed from those calculated in Jain et al. (2000) by more

than 20% (i.e., for CH3Br and CF4). The significant differences between Sihra et al. (2001) and Jain et al. (2000)

for these and some other gases (e.g., HCFC-225cb, HFC134a, HFC-227ea, HFC-23, and HFC-32) suggest that further study and understanding is required before adopting

new forcing efficiency values for these compounds.

SOURCE GASES

1.31

GWP TH 0

TH

0

x TH

x

r

a x t dt

a r t dt

( )

( )

( )

=

[ ]

[ ]

Ú

Ú
SOURCE GASES

1.32

Table 1-6. Direct Global Warming Potentials (mass basis) for gases that have adequately characterized

lifetimes.

Industrial Designation Chemical Radiative Lifetime Global Warming Potential for

or Common Name Formula Efficiency (years) Given Time Horizon

(W m–2 ppb–1) 20 100 500

years years years

Carbon dioxide CO2 1.53 ¥ 10–5 a b 1 11

Methane CH4 3.7 ¥ 10–4 12.0 c 63.

d 23.d 7.d

Nitrous oxide N2O 3.1 ¥ 10–3 114 c 278 300 158

Chlorofluorocarbons

CFC-11 CCl3F 0.25 45 6330 4680 1630

CFC-12 CCl2F2 0.32 100 10340 10720 5230

CFC-13 CClF3 0.25 640 10160 14190 16520

CFC-113 CCl2FCClF2 0.30 85 6150 6030 2700

CFC-114 CClF2CClF2 0.31 300 7560 9880 8780

CFC-115 CClF2CF3 0.18 1700 4990 7250 10040

Hydrochlorofluorocarbons

HCFC-21 CHCl2F 0.14 e 1.7 498 148 46

HCFC-22 CHClF2 0.20 12.0 4850 1780 552

HCFC-123 CHCl2CF3 0.14 f 1.3 257 76 24

HCFC-124 CHClFCF3 0.22 5.8 1950 599 186

HCFC-141b CH3CCl2F 0.14 9.3 2120 713 222

HCFC-142b CH3CClF2 0.20 17.9 5170 2270 709

HCFC-225ca CHCl2CF2CF3 0.20 g 1.9 404 120 37

HCFC-225cb CHClFCF2CClF2 0.32 5.8 1910 586 182

Hydrofluorocarbons

HFC-23 CHF3 0.16 270 9500 12240 10350

HFC-32 CH2F2 0.09 4.9 1800 543 169

HFC-41 CH3F 0.02 2.4 304 90 28

HFC-125 CHF2CF3 0.23 29 5970 3450 1110

HFC-134 CHF2CHF2 0.18 9.6 3200 1090 337

HFC-134a CH2FCF3 0.15 14.0 3370 1320 410

HFC-143 CH2FCHF2 0.13 3.5 1160 347 108

HFC-143a CH3CF3 0.13 52 5540 4400 1600

HFC-152 CH2FCH2F 0.09 0.60 176 52 16

HFC-152a CH3CHF2 0.09 1.4 411 122 38

HFC-227ea CF3CHFCF3 0.30 34.2 5760 3660 1200

HFC-236cb CH2FCF2CF3 0.23 13.6 3420 1320 409

HFC-236ea CHF2CHFCF3 0.30 10.7 3850 1350 420

HFC-236fa CF3CH2CF3 0.28 240 7620 9650 7700

HFC-245ca CH2FCF2CHF2 0.23 6.2 2200 682 212

HFC-245fa CHF2CH2CF3 0.28 7.6 3180 1020 316

HFC-365mfc CH3CF2CH2CF3 0.21 8.6 2370 782 243

HFC-43-10mee CF3CHFCHFCF2CF3 0.40 15.9 3890 1610 502

Chlorocarbons

Methyl chloroform CH3CCl3 0.06 5.0 476 144 45

Carbon tetrachloride CCl4 0.13 26 2540 1380 437

Methyl chloride CH3Cl 0.01 1.3 56 17 5

Bromocarbons

Methyl bromide CH3Br 0.01 0.7 16 5 1
SOURCE GASES

Table 1-6, continued.

Industrial Designation Chemical Radiative Lifetime Global Warming Potential for

or Common Name Formula Efficiency (years) Given Time Horizon

(Wm–2 ppb–1) 20 100 500

years years years

Bromodifluoromethane CHBrF2 0.14 5.8 1290 397 123

Halon-1211 CBrClF2 0.30 16 4460 1860 578

Halon-1301 CBrF3 0.32 65 7970 7030 2780

Halon-2402 CBrF2CBrF2 0.33 h 20 3460 1620 505

Fully fluorinated species

Sulfur hexafluoride SF6 0.52 3200 15290 22450 32780

Trifluoromethylsulfurpentafluoride SF5CF3 0.57 800 12370 17500 21500

FC-14 CF4 0.08 50000 3920 5820 9000

FC-116 C2F6 0.26 10000 8110 12010 18280

FC-218 C3F8 0.26 2600 5940 8690 12520

FC-31-10 C4F10 0.33 2600 5950 8710 12550

FC-318 c-C4F8 0.32 3200 6870 10090 14740

FC-41-12 C5F12 0.41 4100 6120 9010 13330

FC-51-14 C6F14 0.49 3200 6230 9140 13350

Halogenated alcohols and ethers

(CF3)2CFOCH3 0.31 3.4 1130 338 105

(CF3)2CHOH 0.28 2.0 718 214 66

HFE-125 CHF2OCF3 0.44 136 12970 14670 8530

HFE-134 CHF2OCHF2 0.45 26 11470 6220 1970

HFE-143a CH3OCF3 0.27 4.3 2480 744 231

HFCE-235da2 CHF2OCHClCF3 0.38 2.6 1150 343 107

HFE-245cb2 CH3OCF2CF3 0.32 5.1 2300 697 216

HFE-245fa2 CHF2OCH2CF3 0.31 4.9 2150 649 201

HFE-254cb2 CH3OCF2CHF2 0.28 2.6 1190 353 110

HFE-347mcc3 CH3OCF2CF2CF3 0.34 5.2 1860 566 176

HFE-356pcf3 CHF2OCH2CF2CHF2 0.39 3.6 1660 494 153

HFE-374pc2 i CH3CH2OCF2CHF2 0.25 5 1810 548 170

HFE-7100 C4F9OCH3 0.31 5 1310 397 123

HFE-7200 C4F9OC2H5 0.30 0.77 189 56 17

H-Galden 1040x j CHF2OCF2OC2F4OCHF2 1.37 6.3 5940 1840 572

HFE-236ca12 CHF2OCF2OCHF2 0.66 12.1 7560 2780 864

HFE-338pcc13 CHF2OCF2CF2OCHF2 0.87 6.2 4770 1480 459

Lifetimes are from Table 1-3. GWPs for some gases that were tabulated in IPCC (2001) are not reported here because they have lifetimes less than 0.5

years (see notes to Table 1-3).

a Assumes a global surface mean mixing ratio of 370 ppm.

b CO2 atmospheric response function is as described in WMO (1999) and Ramaswamy et al. (2001). c These lifetimes are adjustment times including feedbacks of emissions on lifetimes (see Ramaswamy et al., 2001).

d The GWPs listed for methane include indirect effects of tropospheric ozone production and stratospheric water vapor production (Solomon and

Wuebbles et al., 1995; Ramaswamy et al., 2001). e The Sihra et al., (2001) radiative efficiency for HCFC-21 of 0.14 W m–2 ppb–1 is adopted rather than the 0.17 W m–2 ppb–1 value in WMO (1999)

and Ramaswamy et al. (2001). The value of 0.17 was incorrectly tabulated in WMO (1999) (see Sihra et al. (2001) for discussion of error).

f The radiative efficiency from Sihra et al. (2001) and Jain et al. (2000) was adopted here because these studies were in good agreement and because

the estimate adopted by Ramaswamy et al. (2001) originated from Fisher et al. (1990), in which a lower resolution calculation was performed. g The radiative efficiency from Sihra et al. (2001) and Jain et al. (2000) was adopted here because these studies were in good agreement and because

the estimate adopted by Ramaswamy et al. (2001) was based upon an unpublished calculation made for WMO (1995).

h Radiative efficiency from Sihra et al. (2001).

i Referred to as HFE-374pcf2 in past Assessments; however, the current name unambiguously defines the compound.

j Also known as HFE-43-10pccc124.

1.33
1.4.5.2 DIRECT GLOBAL WARMING POTENTIALS

The “direct” portion of a compound’s GWP represents the part that is due to the infrared absorptive and

emissive properties that lead to a reduction in the net radiation exiting the troposphere. Here we have divided our

compilation of direct GWPs into two tables. Table 1-6

includes compounds whose lifetimes are relatively well

established, whereas Table 1-7 has compounds whose lifetimes are more uncertain, either because of a dearth of

laboratory measurements or an incomplete understanding

of the loss processes. Compounds characterized by lifetimes less than 0.5 years are not included in the tables

because of the difficulties involved in making meaningful

estimates of a global lifetime and a global radiative effect.

In comparing Tables 1-6 and 1-7 with past direct GWP

tables (e.g., Granier and Shine et al., 1999; Ramaswamy

et al., 2001), note that there is now an additional significant figure tabulated for all GWP values greater than 10.

This does not imply that there is any improvement in

the accuracy of the GWP values, which is believed to be

±35%. The additional significant figure is included to

facilitate tracking GWP changes in the future by reducing

the impact of rounding.

1.4.5.3 NET GLOBAL WARMING POTENTIALS

As discussed in past Assessments, the impacts of

many chlorocarbons and bromocarbons on climate are

complicated by their destruction of stratospheric ozone.

The destruction of ozone offsets part, and in some cases

all, of the direct forcing for the period in the future when

ozone destruction occurs. Table 1-8 includes estimates of

these net GWPs for selected halocarbons. These values

are updated from Daniel et al. (1995) for revised halocarbon lifetimes (Table 1-3), assumption of the relative

effectiveness for ozone destruction of bromine compared

with chlorine (assumed to be 45 here), future ozone loss

scenario (using EESC from scenario Ab; see Section 1.8),

and time of emission (assumed in year 2002 here).

The net GWP values here differ from those in

Ramaswamy et al. (2001) primarily because of the different assumed time of emission of the radiatively active

compound in conjunction with the time dependence of

polar and global ozone depletion.

1.4.6 Hydroxyl Radical Changes with Time

Reaction with tropospheric hydroxyl radical (OH)

is the chief sink for many of the halocarbons in the comSOURCE GASES

1.34

Table 1-7. Direct Global Warming Potentials (mass basis) relative to carbon dioxide for gases whose

lifetimes are determined only by indirect means rather than by laboratory measurements, or for which

uncertainty exists in our understanding of the important loss processes.

Industrial Chemical Radiative Lifetime Global Warming Potential for

Designation or Formula Efficiency (years)Given Time Horizon

Common Name (W m–2 ppb–1)20 100 500

years years years

Nitrogen trifluoride NF3 0.13 740 7780 10970 13240

Perfluorocyclopropane c-C3F6 0.42 >1000 >11950 >17070 >21920

HFE-227ea CF3CHFOCF3 0.40 11.0 4270 1520 471

HFE-236ea2 CHF2OCHFCF3 0.44 5.8 3170 973 302

HFE-236fa CF3CH2OCF3 0.34 3.7 1610 480 149

HFE-245fa1 CHF2CH2OCF3 0.30 2.2 948 282 88

HFE-329mcc2 CF3CF2OCF2CHF2 0.49 6.8 2880 904 281

HFE-338mcf2 CF3CF2OCH2CF3 0.43 4.3 1810 543 169

HFE-347mcf2 CF3CF2OCH2CHF2 0.41 2.8 1240 368 114

HFE-356mec3 CH3OCF2CHFCF3 0.30 0.94 334 99 31

HFE-356pcc3 CH3OCF2CF2CHF2 0.33 0.93 363 108 34

HFE-356pcf2 CHF2CH2OCF2CHF2 0.37 2.0 876 260 81

(CF3)2CHOCHF2 0.41 3.1 1250 373 116

-(CF2)4CH(OH)- 0.30 0.85 239 71 22

All lifetimes were taken from Ramaswamy et al. (2001). The GWPs differ from those in Ramaswamy et al. (2001) because an additional significant

figure was retained and because an updated CO2 mixing ratio was assumed (370 ppm).
pendium of ODPs and GWPs discussed in this chapter.

The halocarbon lifetimes, on which the ODPs and GWPs

depend, are determined by the morphology of the steadystate halocarbon distributions, the kinetics of the OH reaction, and the OH distribution. The time horizon for halocarbon effects summarized by ODPs and GWPs is decades

to centuries, and halocarbon lifetimes with respect to tropospheric losses can be decades. It is noteworthy that

time invariance of the lifetimes is a typical assumption

made to simplify the explanation of these concepts, which

also effectively implies an unchanging OH distribution.

Karlsdóttir and Isaksen (2000), however, present a multidimensional model result for the period 1980-1996 that

produces an increase in mean global tropospheric OH

levels of 7% over the period, driven largely by increases

in low-latitude emissions of nitrogen oxides (NOx) and

carbon monoxide (CO).

The distribution of OH is controlled by local photochemistry comprising the interactions of ozone, water,

solar ultraviolet radiation, CO, methane (CH4), nitrogen

oxides, and higher hydrocarbons. Wang and Jacob (1998)

describe the relationship of OH abundance to a ratio of

nitrogen oxide and hydrocarbon concentrations and

ascribe the apparent constancy of OH over the historical

industrial period to the conservation of this ratio even as

emissions for both increased. The model and analysis of

Wang and Jacob (1998) suggests a 9% decline in weighted

global OH from preindustrial time, but Prather and Ehhalt

et al. (2001) note that no consensus exists on the magnitude of this change.

The importance of the OH distribution and its

changes with time in the context of this chapter includes

the effect of interannual variability on the methyl chloroform (CH3CCl3) reference OH lifetime. On longer time

scales, changes in the OH distribution would change some

halocarbon lifetimes, ODPs, and GWPs, increasing or

decreasing their impacts relative to the current values.

Future increases in direct and indirect greenhouse gases

could produce these changes through direct participation

in OH-controlling chemistry, indirectly through stratospheric ozone changes that could increase solar ultraviolet in the troposphere, and potentially through climate

change effects on biogenic emissions, humidity, and

clouds. Prinn et al. (2001) analyzed a 22-year record of

global CH3CCl3 measurements and emission estimates

and suggest that global OH over that period varied from

about 5% above to 10% below the period average. This

inferred variability in OH contributes to the uncertainty

of the reference CH3CCl3 lifetime and other lifetimes

based on it. These results through 1993 are essentially

consistent with the conclusions of Krol et al. (1998, 2001),

who used the same measurement and emission record but

SOURCE GASES

1.35

Table 1-8. Net Global Warming Potentials of selected halocarbons (on an equal mass basis).

Time Horizon = 2022 Time Horizon = 2102

(20 years) (100 years)

Direct Net Direct Net

Species Min Max Min Max

CFC-11 6330 –938 4910 4680 –524 3660

CFC-12 10340 6660 9620 10720 7660 10120

CFC-113 6150 1770 5290 6030 2470 5330

HCFC-22 4850 4020 4690 1780 1420 1710

HCFC-123 257 –147 177 76 –44 53

HCFC-124 1950 1490 1860 599 450 570

HCFC-141b 2120 –34 1700 713 –109 552

HCFC-142b 5170 4260 4990 2270 1810 2180

CH3CCl3 476 –2080 –26 144 –659 –13

CCl4 2540 –5460 972 1380 –3440 433

CH3Br 16 –7650 –1490 5 –2270 –444

Halon-1211 4460 –67410 –9630 1860 –33360 –5050

Halon-1301 7970 –77490 –8780 7030 –59280 –5970

Net Global Warming Potentials are calculated from updated lifetimes and radiative data reported in this chapter, and include the indirect cooling effects

that result from stratospheric ozone depletion as well as the direct radiative effects. The magnitude of the cooling is estimated using the future EESC

of scenario Ab, a value for alpha of 45, and an emission pulse in the year 2002. The min/max ranges of the net GWPs reflect the uncertainty in the

change in ozone radiative forcing from 1980 to 1990 of –0.153 to –0.03 W m–2 (Ramaswamy et al., 2001).
an independent calculation technique to infer an increase

of 0.46% yr-1 over the shorter period of 1978-1993. Prinn

et al. (2001) infer that OH values in the late 1990s are

below those in the late 1970s to early 1980s, in agreement

with the longer CH3CCl3 lifetime reported by Montzka et

al. (2000) for 1998-1999 relative to the Prinn et al. (2001)

value for the full-period average. Inferences regarding

methyl chloroform lifetimes or trends in OH are quite sensitive to any errors in the absolute magnitude of estimated

emissions. These errors could be significant and enhanced

during the late 1990s because annual emissions were dropping precipitously then. Independent estimates of methyl

chloroform emissions and their changes during the late

1990s would provide a useful independent assessment of

these implied trends in global OH.

IPCC (Prather and Ehhalt et al., 2001), using scenarios generated by the IPCC Special Report on Emissions

Scenarios (Nakic ́enovic ́ et al., 2000) and on the basis of a

comparison of results from 14 models, predicts that global

OH could vary from a 20% decrease to a 6% increase by

2100. The range of these predicted changes is not materially greater than that already calculated by Prinn et al.

(2001). On the other hand, a 9% decrease from preindustrial time to the present and a projected 20% decrease

present to future would, in combination, give a substantial change from the preindustrial atmospheric state. The

sense of the change is indicated to be most likely in the

direction of longer halocarbon lifetimes with respect to

oxidation by OH and enhanced atmospheric effects.

1.4.7 Other Sinks

Atmospheric in situ halocarbon sinks include photolysis and homogeneous gas-phase reactions with OH,

Cl, and O(1D). The dynamics of tropospheric mixing

exposes tropospheric air to contact with the surface and

allows for irreversible deposition as well. Organisms in

surface ocean waters can both consume and produce halocarbons. Chemical degradation of dissolved halocarbons

can also occur through hydrolysis. Physical dissolution

into ocean waters occurs, but does not represent a significant sink for these halocarbons. These processes are

highly variable in the ocean, depending on physical

processes of the ocean mixed layer, temperature,

productivity, surface saturation, and other variables.

Determining a net global sink through observation is a

difficult task. Yvon-Lewis and Butler (2002) have constructed a high-resolution model of the ocean surface

layer, its interaction with the atmosphere, and physical,

chemical, and biological ocean processes. They examined ocean uptake for a range of halocarbons, from

knowledge of solubilities and chemical and biological

degradation rates. Although they do not find the ocean

sink to be significant for HFCs and HCFCs compared with

in situ atmospheric sinks, the reported 94-year oceanic

loss lifetimes for CH3CCl3 and, coincidentally, for CCl4

as well, must be included in determining the total lifetimes of these compounds.

Particular consideration of these surface sinks is

pertinent for the halocarbons that serve as reference compounds, e.g., CFC-11 and CH3CCl3, and for halocarbons

with negligible atmospheric sinks. Kutsuna et al. (2000)

have observed degradation of CH3CCl3 on the surface of

aluminosilica clay mineral surfaces under laboratory conditions. They conclude that both on suspended particles

under dust-storm conditions and on surface material in

deserts, local sink rates may grow large enough to come

within an order of magnitude of atmospheric sinks.

Estimates of sink magnitudes on a global basis are not yet

available, but would appear to be small.

1.5 METHYL BROMIDE AND 

METHYL CHLORIDE

Methyl bromide and methyl chloride differ from

the other “long-lived” ozone-depleting gases because they

have significant natural and anthropogenic sources.

Recent research has focused on identifying and quantifying the natural sources and sinks of these two compounds. The calculated atmospheric budgets of both

gases are largely out of balance, with identified sinks outweighing identified sources.

1.5.1 Methyl Bromide (CH3Br)

Since the 1998 Scientific Assessment (WMO,

1999), many publications have appeared and have

addressed the atmospheric history of methyl bromide

(CH3Br), its concentration and distribution in the atmosphere (see Section 1.2.1.6 of this chapter), its oceanic and

terrestrial fluxes, and its mechanisms of production and

degradation in nature. These studies have allowed for

refinements in our understanding of atmospheric methyl

bromide, but many of the significant uncertainties in 1998

remain. Although new sources have been identified, they

do not balance current estimates of loss, which remain

essentially unchanged from the previous Assessment

(Table 1-9). Analyses of firn air provide for the first time

an inferred, 20th century trend for atmospheric methyl bromide in the Southern Hemisphere. These results may provide additional constraints to our understanding of source

and sink magnitudes. Interlaboratory differences for

absolute calibration of atmospheric methyl bromide are

approximately 10%.

SOURCE GASES

1.36
1.5.1.1 ANTHROPOGENIC AND ANTHROPOGENICALLY

INFLUENCED SOURCES

The main anthropogenic sources of CH3Br include

fumigation of soils for planting and of agricultural products for import and export, biomass burning inasmuch as

it is influenced by humans, and automobile emissions

from burning of leaded gasoline (Table 1-9). Possible

alterations of natural emissions associated with changes

in land-use patterns are not included here, nor are there

data available to quantify them.

Methyl bromide use as a fumigant in non-Article

5(1) countries (developed countries as defined by the

Montreal Protocol) was scheduled to have begun

decreasing after 1998. The Montreal Protocol specifies

that production of CH3Br in developed countries be

reduced from the 1991 levels by 25% in 1999 and by 50%

in 2001. Data reported to the UNEP Ozone Secretariat

SOURCE GASES

1.37

Table 1-9. Summary of the estimated source and sink strengths of methyl bromide (CH3Br).

Source or Sink Type Source or Sink Best Estimate Source or Sink Full Range 

(Gg yr–1) (Gg yr–1)

Sources

Ocean a 63 23 to 119

Fumigation–soils b 26.5 16 to 48

Fumigation–durables b 6.6 4.8 to 8.4

Fumigation–perishables b 5.7 5.4 to 6.0

Fumigation–structures b 2 2 to 2

Gasoline 5 0 to 10

Biomass burning 20 10 to 40

Wetlands* 4.6 2.3 to 9.2

Salt marshes* 14 7 to 29

Shrublands* 1 0.5 to 2

Rapeseed* 6.6 4.8 to 8.4

Rice fields* 1.5 0.5 to 2.5

Fungus* 1.7 0.5 to 5.2

Peatlands* 0.9 0.1 to 3.3

Subtotal (Sources) 159 77 to 293

Sinks

Ocean a –77 –37 to –133

OH and hn –80 –60 to –100

Soils c –47 –32 to –154

Plants (not quantified) (not quantified)

Subtotal (Sinks) –204 –129 to –387

Total (Sources + Sinks) d –45 –220 to +71

See text for further explanation and references related to the values listed. Ranges listed for wetlands and shrublands were assumed here to be a factor

of 2. Loss due to hydroxyl has been revised, considering the updated analysis of methyl chloroform (see Section 1.4). Table is adapted and augmented from Yvon-Lewis (2000).

* All asterisked items were estimated from measurements of net fluxes and may be influenced by sinks within them. Thus, they represent minimum

gross fluxes.

a The oceanic source has been adjusted to accommodate the slightly smaller undersaturations reported since the 1998 Assessment (WMO, 1999; see 

text). The source is calculated as source = net oceanic flux – sink, where net oceanic flux is –14 (–11 to –20) Gg yr–1. b Fumigation estimates are taken directly from Kurylo and Rodríguez et al. (1999) and are based upon industrial production data for 1992. c Revised uptake to soils is included here as discussed in Varner et al. (1999a), Crill (2000), and Yvon-Lewis (2000).

d The possible range for sources plus sinks here is not a straightforward calculation of highest and lowest values, because the oceanic net flux must

hold to within the range of –11 to –20 Gg yr–1. See p. 2.22 in Kurylo and Rodríguez et al. (1999) for a complete explanation. The total range given

in this table is a range of extremes, requiring all sinks to be at their highest values while all sources are at their lowest possible values and vice versa.

The likelihood of such occurrences is extremely low. Root mean square error on total sinks of –204 Gg yr–1 is estimated from the indicated full

ranges as –157 to –327 Gg yr–1; root mean square error of –45 Gg yr–1 on the sum of sources and sinks is estimated at -171 to +14 Gg yr–1.
(UNEP, 2002) are consistent with such a reduction and

suggest further that global consumption in 1999 was ~11

Gg less than that in 1998.

Little has changed in our understanding of the

degree to which the combustion of leaded fuel contributes

to emissions of atmospheric CH3Br. Currently about 86%

of all gasoline sold in the world is unleaded, and this will

likely rise to 89% by 2005 (personal communication, M.P.

Walsh, international consultant, Arlington, Virginia, U.S.,

2001). Automobile emissions today remain a small part

of the global budget (5 ± 5 Gg yr-1, or ~3% of the total

emissive flux estimated from sinks), although they were

substantially larger in the past (Thomas et al., 1997).

Similarly, little new information on the emission of methyl

bromide from biomass burning has emerged since 1998.

The studies of N.J. Blake et al. (1996) and Mano and

Andreae (1994) reported in the 1998 Assessment (WMO,

1999) remain our best estimates, supporting globally averaged emissions of 20 (10-40) Gg yr-1

.

1.5.1.2 SURFACE INTERACTIONS: NATURAL SOURCES

AND SINKS

Ocean

New work continues to confirm that the ocean

remains the largest identified source and second largest

sink for atmospheric CH3Br. This has allowed for refinements in our understanding of methyl bromide saturation,

and there is some suggestion that methyl bromide in temperate waters is supersaturated in the spring and summer

and undersaturated in the fall and winter (King et al.,

2000), which is consistent with observations of seasonal

cycles in the North Sea by Baker et al. (1999). This could

explain some of the apparently anomalous supersaturations in the open ocean observed in previous investigations (e.g., Lobert et al., 1996; Groszko and Moore, 1998),

although not all observed saturation anomalies fit such a

pattern (e.g., King et al., 2000; Northeastern Atlantic data

in Baker et al.,1999). Combining results from the various

expeditions, and using a bimodal relationship to temperature similar to that suggested by Groszko and Moore

(1998), King et al. (2000) determined that the overall

effect of including these seasonal differences on a global

scale lowered the estimated net flux slightly downward

from –21 (–32 to –3) to –14 (–20 to –11) Gg yr-1

. These

findings do not affect the estimate of lifetime relative to

oceanic loss, but they do reduce the budget imbalance

slightly (lifetime relative to oceanic loss is calculated

directly from rates of hydrolysis and nucleophilic substitution in seawater and rates of air-sea exchange).

Two isotopic studies of CH3Br consumption in the

open ocean have been published, supporting the earlier

finding of significant biological removal of CH3Br in seawater (Tokarczyk and Saltzman, 2001; Tokarczyk et al.,

2001). The best estimate of the partial lifetime of atmospheric CH3Br with respect to oceanic loss remains 1.9

(1.1-3.9) years (Yvon-Lewis and Butler, 1997), as reported

in the 1998 Assessment (WMO, 1999).

Soils

Soil fluxes of CH3Br are slightly better constrained

than in 1998. The soil sink reported in the 1998

Assessment (WMO, 1999) was large, –42 Gg yr-1, but

also included large uncertainties (–214 to –10 Gg yr-1).

This wide range stems from two studies reporting substantially different means. Shorter et al. (1995) had estimated a global sink of –42 ± 32 Gg yr–1, whereas Serca et

al. (1998) arrived at –143 ± 70 Gg yr-1. The difference

between these two studies was in part due to the measurements for cultivated soils, but was also a function of the

choice of procedure for extrapolating. The estimate for

losses to agricultural soils is much larger in the Serca et

al. (1998) study. Because of this, Serca et al. (1998) also

suggested a mean flux of –94 ± 54 Gg yr-1 by taking an

arithmetic mean of the results from the two studies for

agricultural soils. After making additional measurements

at 40 globally distributed sites, Varner et al. (1999a)

revised their estimate of agricultural losses upward, but

only by a small amount. Based upon these new data and

considerations of previously reported data, Crill (2000)

revised the range for loss of atmospheric methyl bromide

to soils, narrowing it to –154 to –32 Gg yr-1. Adjusting

the total soil loss for the new measurements of losses to

cultivated soils, Yvon-Lewis (2000) suggested a mean flux

of –47 Gg yr-1, within the range reported by Crill (2000).

Recent reports have suggested that at least some

soils can produce CH3Br and even emit CH3Br to the

atmosphere. Keppler et al. (2000) identified an abiotic

mechanism for the production of halogenated hydrocarbons in soil that yields high production rates. This mechanism involves iron, halide ion, and organic matter, and

the reaction can be fast. How much of this production

reaches the atmosphere, and whether such production is

limited to peatlands or is ubiquitous, is presently

unknown. The study of Rhew et al. (2001) showed that

shrublands could be a net source or sink of CH3Br and

CH3Cl. Because this study focused on net fluxes, it does

not relate directly to estimates of lifetime.

Plants and Plant Ecosystems

In the 1998 Assessment (WMO, 1999), a potentially significant plant sink was reported but not quantified globally (Jeffers and Wolfe, 1997). Because of a lack

SOURCE GASES

1.38
of additional studies related to this loss, calculating a partial lifetime of atmospheric methyl bromide from plants

is not yet possible.

In the past 4 years, however, considerable attention

has been given to the study of net fluxes from isolated,

terrestrial ecosystems. Dimmer et al. (2001) measured

CH3Br net fluxes from Irish peatlands, Varner et al.

(1999b) measured net fluxes from New Hampshire wetlands, and Rhew et al. (2000, 2001) studied the net fluxes

of methyl bromide from salt marshes and shrublands.

Extrapolating these measurements globally yields net flux

estimates of 0.9 Gg yr-1 for peatlands, 4.6 Gg yr-1 for wetlands, and 14 Gg yr-1 for salt marshes. Shrublands were

reported as having a net flux of <1 Gg yr-1. Lee-Taylor

and Holland (2000), modeled published fungal production rates, deriving a global net flux of 1.7 Gg yr-1

. From

studying cultivated crops, Gan et al. (1998) calculated a

net flux of 6.6 Gg yr-1 for rapeseed, and Redeker et al.

(2000) calculated a net flux of 1.5 Gg yr-1 for rice fields.

What is important to note is that all these studies report

measurements of net fluxes and, although useful in narrowing the gap between sources and sinks in the atmospheric budget of CH3Br, they are not suited alone for determining gross fluxes or partial atmospheric lifetimes.

1.5.1.3 ATMOSPHERIC LIFETIME

Although new publications have confirmed biological losses in the ocean, refined the magnitude of losses to

soils, and revised our understanding of global OH burdens (see Section 1.4.2), the best estimate of the methyl

bromide global lifetime remains 0.7 years (0.5-0.9 year

root mean square (rms) uncertainty; 0.4-1.1 year full

range; Kurylo and Rodríguez et al. (1999)). A couple of

things have the potential to change the calculated atmospheric lifetime of CH3Br, however. First, the magnitude

of the soil sink remains highly uncertain. Measurements

have been made that help in reducing uncertainties, but

the result also depends upon the method of extrapolation

chosen. Furthermore, the determination that any additional sinks, such as plants, might be significant would

reduce further the estimated lifetime for methyl bromide.

1.5.1.4 A TWENTIETH CENTURY, SOUTHERN

HEMISPHERIC TREND AND ITS IMPLICATIONS

Since the previous Assessment (WMO, 1999), two

papers have inferred atmospheric histories for methyl bromide from measurements of Antarctic firn air (Butler et

al., 1999; Sturges et al., 2001a). These studies, based upon

samples collected at a total of four locations in Antarctica,

suggest that mixing ratios of methyl bromide in the

Southern Hemisphere have increased by about 3 ppt since

1900 (Figure 1-11). The 2-ppt increase suggested for the

latter half of the last century is roughly coincident with

the onset of methyl bromide use as a fumigant in the early

1960s (Singh and Kanakidou, 1993), but other sources or

sinks also may have changed over that time as well. The

rate of increase of about 0.6% yr-1 from about 1970 to

1990 inferred from firn air agrees with the real-time measurements of Khalil et al. (1993) for the Southern

Hemisphere during 1978-1993 and with the data obtained

from archived flask air filled at Cape Grim, Tasmania,

over this same period (Figure 1-11; Miller, 1998). The

SOURCE GASES

1.39

4

5

6

7

8

9

1 0

1885 1905 1925 1945 1965 1985 2005 C H3BrMixingRatio(ppt) Figure 1-11. Southern Hemispheric history of

methyl bromide suggested by Antarctic firn air measurements from four locations, and by archived air

samples and flask samples collected at Cape Grim,

Tasmania (41°S). The solid line is an atmospheric

history derived from firn air measurements at South

Pole, Antarctica, estimated with a vertical diffusion

model that fits the history of the CH3Br/CO2 ratio to

the observed CO2 and CH3Br firn air profiles (see

Butler et al., 1999, for details). The green circles

are firn data from Siple Dome, Antarctica, plotted

against dates determined from CO2 measurements

and the ratio of CH3Br and CO2 diffusivities (Butler

et al., 1999); the red circles are from firn air at

Dronning Maud Land, Antarctica, and are plotted

against dates adjusted from CFC-12 measurements

and the ratio of CH3Br and CFC-12 diffusivities

(Sturges et al., 2001a); blue diamonds are similar

data for Dome C, Antarctica (Sturges et al., 2001a);

and yellow (Khalil et al., 1993) and light blue (Miller,

1998) plus symbols are from analyses of flask samples collected at Cape Grim.
work of Butler et al. (1999), which reported results for

firn air dating back to the end of the 19th century, indicated that atmospheric methyl bromide may have

increased as much as 1 ppt in the first half of the 20th century, but the causes for this earlier apparent increase, and

whether they continued into the latter half of the century,

are unknown.

Atmospheric histories derived from the analysis of

firn air, however, have significant uncertainties. Dating

of firn air is imprecise (Schwander et al., 1988), and the

atmospheric histories derived for methyl bromide presume that in situ production or degradation are insignificant during the long time that the firn air is exposed to the

surrounding snow. Whereas all Antarctic sites sampled

have yielded a similar atmospheric history for CH3Br

regardless of their differences in temperature, deposition

rate, depth, and lock-in zone thickness (Figure 1-11),

Arctic sites exhibited large elevations of methyl bromide

and other gases near the bottoms of the profiles (not

shown; Butler et al., 1999; Sturges et al., 2001a). Further,

the profiles from the Northern Hemispheric sites differed

dramatically from one another, as would be expected for

site-specific disturbances. That this was not observed in

the Antarctic increases confidence in the reported atmospheric history from these sites. The mechanism for

increases of certain compounds, including methyl bromide, at the bottom of the Arctic profiles is not known.

An understanding of atmospheric changes for

methyl bromide during the 20th century can provide some

insight into the relative magnitude of present-day sources.

Of particular interest is the magnitude of industrially

derived emissions relative to the total annual flux of

methyl bromide. This quantity can be estimated directly

from our understanding of source and sink magnitudes:

for emissions derived from fumigation of 41 Gg yr-1, with

an rms range of 30-62 Gg yr-1, divided by the best estimates for sinks (204 Gg yr-1 and rms range of 157-327;

Table 1-9), the calculated contribution of industrially produced methyl bromide to a 10-ppt global atmospheric

burden ranges from 10 to 40%. As suggested above, the

atmospheric trend inferred from the Antarctic firn air provides a second approach for estimating this fraction. The

results suggest that from 1950 to 1995, CH3Br in the

Southern Hemisphere increased by 2 to 2.5 ppt (Figure

1-11). If we assume that either Northern Hemispheric

mixing ratios increased by the same proportion as in the

Southern Hemisphere or that Northern Hemispheric

mixing ratios in 1950 were the same as those in the

Southern Hemisphere, then the firn results suggest that

CH3Br in the global atmosphere has increased 25-40%

from 1950 to 1995, regardless of lifetime or other budget

considerations. Therefore, were this increase solely due

to fumigation, 40% would represent an upper limit for the

contribution from fumigation. However, it also appears

from the firn data that CH3Br in the Southern Hemisphere

increased about 0.7 ppt during the first half of the century,

perhaps because of changing land-use practices, increases

in biomass burning, or the post-1925 onset of leaded-fuel

combustion in automobiles. If this underlying trend had

continued through the second half of the century, then the

contribution of industrially produced methyl bromide

would be 1.5-3 ppt, or 15-30% of the 1995 atmospheric

burden. Given the assumptions above, then, the overall

range of 15-40% for this fraction suggests emissions

related to fumigation of 30-80 Gg yr-1 in 1995. Although

bromine emissions from combustion of leaded fuel have

decreased dramatically since 1970, they were similar in

1950 to what they were in 1990 (Thomas et al., 1997) and

do not strongly influence this conclusion.

1.5.1.5 ATMOSPHERIC BUDGET

The calculated budget for atmospheric methyl bromide remains out of balance, with estimated sinks still

outweighing estimated sources (Table 1-9). However, the

new findings on ecosystem net fluxes have narrowed the

gap between calculated sources and sinks (Crill, 2000;

Yvon-Lewis, 2000; Butler, 2000). The most complete

summary is that of Yvon-Lewis (2000), from which Table

1-9 has been adapted and augmented. This budget is for

the mid-1990s, before the scheduled reductions of methyl

bromide production. Because measurements relating to

plants usually involve estimates of net fluxes, any plant

sink is incorporated implicitly into the source terms for

those elements that are reported as such in the table.

Quantifying the plant sink, while affecting the calculated

lifetime, would not affect the budget balance.

1.5.2 Methyl Chloride (CH3Cl)

Methyl chloride (CH3Cl), the most abundant halocarbon in the atmosphere, is derived largely from natural

sources. Until 1996 most of the input to the atmosphere

was considered to originate from the oceans, but investigations in recent years have indicated that terrestrial

sources, probably located in the tropics, dominate the

atmospheric budget, and that these sources are more significant than hitherto appreciated. Moreover, the presumption that abiotic reactions constitute the only sinks

for atmospheric CH3Cl is being displaced by recognition

that microbial degradation in soil may represent a measurable sink of atmospheric CH3Cl. Although new findings have allowed constraints to be placed on some source

SOURCE GASES

1.40
terms in the atmospheric budget, major uncertainties still

exist regarding the magnitude of several sources and sinks.

As is seen for methyl bromide, a shortfall is apparent

between known sources and modeled sinks of methyl

chloride.

1.5.2.1 ATMOSPHERIC DISTRIBUTION AND TRENDS

There is emerging evidence that CH3Cl is elevated

in air masses associated with land in the tropics (Khalil

and Rasmussen, 1999; Yokouchi et al., 2000b; see also

Section 1.2.1.7). CH3Cl mixing ratios of up to 1400 ppt

have been recorded in calm weather off the subtropical

Okinawa and Hateruma Islands in southern Japan (Li et

al., 1999; Yokouchi et al., 2000b) and up to 850 ppt at

inland locations on various continents (Khalil and

Rasmussen, 1999). Interestingly, the high CH3Cl mixing

ratios observed by the Japanese researchers were strongly

correlated with elevated amounts of the plant-derived terpene, a-pinene, which has an atmospheric lifetime of only

a few hours. This finding is consistent with high emissions of CH3Cl from tropical coastlands. In subsequent

work on the marine boundary layer in the Western Pacific

and Southeastern Indian Ocean, Li et al. (2001) found that

the CH3Cl mixing ratio was enhanced (up to 950 ppt) in

air masses derived from forested tropical islands.

An indication of long-term trends in atmospheric

CH3Cl has been obtained by analysis of air trapped in

polar firn in Antarctica and Greenland (Figure 1-12; Butler

et al., 1999). An increase of about 10% was inferred since

the mid-1900s for both hemispheres, suggesting that the

anthropogenic contribution, if reflected by the apparent

20th century increase, is comparatively small. Global

measurements from 1981-1997 suggest an overall

decrease of about 4% over this shorter period (Khalil and

Rasmussen, 1999).

1.5.2.2 SOURCES OF ATMOSPHERIC METHYL

CHLORIDE

The principal sources of atmospheric CH3Cl identified to date are biomass burning and the oceans, although

tropical plants may exceed all other sources (Yokouchi et

al., 2002; Table 1-10). Except for the tropical plant source,

the latest estimates of emissions do not differ substantially

from those reported in the 1998 Assessment (WMO, 1999;

Graedel and Keene, 1999). The net flux from the oceans

was revised sharply downward in 1996 to 200-400 Gg

yr-1 (Moore et al., 1996). This estimate represents the balance between emissions of CH3Cl in tropical waters and

uptake in the cold waters of latitudes above 50°. Khalil et

al. (1999) analyzed data collected by several groups of

investigators between 1983 and 1996 and confirmed the

weaker oceanic source, calculating a net global flux of

600 Gg yr-1 (after adjusting for calibration). For biomass

burning, Lobert et al. (1999) estimated emissions of

CH3Cl from nine different burning categories on a 1° latitude by 1° longitude grid. Fluxes were derived from biomass inventories based upon emission ratios relative to

both CO and CO2 and the chlorine content of the fuel feedstock. Calculated global emissions of 911 Gg yr-1 (range

655-1125 Gg yr-1) from burning, most of which are

believed confined to the tropics and subtropics, agree with

the previous approximation of 1.0 ± 0.5 Tg yr-1 (Andreae

et al., 1996; N.J. Blake et al., 1996; Rudolph et al., 1995).

Recently, Yokouchi et al. (2002) suggested that

methyl chloride emissions from tropical plants might be

SOURCE GASES

1.41

300

350

400

450

500

550

600

1885 1905 1925 1945 1965 1985 2005 CH3Cl Mixing Ratio (ppt) Figure 1-12. Southern Hemispheric history of

methyl chloride suggested by Antarctic firn air measurements from two locations, and by atmospheric

measurements from flask samples collected in

Antarctica and at Cape Grim, Tasmania. The solid

line is an atmospheric history derived from firn air

measurements at South Pole, Antarctica, estimated

with a vertical diffusion model that fits the history of

the CH3Cl/CO2 ratio to the observed CO2 and CH3Cl

firn air profiles (see Butler et al., 1999, for details).

The open green circles are data from Siple Dome,

Antarctica, plotted against dates derived from

adjusted CO2 measurements and diffusivities (Butler

et al., 1999), and the open blue diamonds and red

triangles are annual means of “real-time” measurements from flask samples collected in Antarctica and

Cape Grim, respectively (Khalil and Rasmussen,

1999; these data have been lowered by 8.3% to

account for a calibration difference).
the largest known source. They determined that a specific group of ferns and trees in Southeast Asia alone produced 910 Gg yr-1

, about one-fourth of the annual flux as

estimated by sinks. The uncertainty in this estimate is

large, and this source certainly warrants additional investigation.

Release of CH3Cl by wood-rotting fungi has been

estimated by Watling and Harper (1998) at 160 Gg yr-1

(Table 1-10), of which 75% is derived from tropical and

subtropical forests. This assessment was based on the

conservative assumption that only Cl- from wood itself

was available for fungal uptake. If, however, Cl– in leaf

litter and soil is accessible to these fungi, emissions could

be much greater. Dimmer et al. (2001) reported field

measurements of CH3Cl fluxes from forest floors.

Although the sites (conifer plantations in Western Ireland)

cannot be regarded as globally representative of temperate

coniferous forest, an extrapolation of the observed emission fluxes led to an estimate of 85 Gg yr-1 from such

forests worldwide, suggesting that the order of magnitude

of the overall flux from fungi in forest ecosystems is similar to that predicted by Watling and Harper (1998).

SOURCE GASES

1.42

Table 1-10. Estimated source and sink strengths for atmospheric methyl chloride (CH3Cl).

Source or Sink Type Source or Sink Best Estimate Source or Sink Full Range

(Gg yr–1) (Gg yr–1)

Sources

Oceans 600 a 325 to 1300

Biomass burning 911 b 655 to 1125

Tropical plants 910 c 820 to 8200

Fungi 160 d 43 to 470

Salt marshes 170 e 65 to 440

Wetlands 40 f, g 6 to 270

Coal combustion 105 h 5 to 205

Incineration 45 h 15 to 75

Industrial 10 h

Rice 5 i

Subtotal (Sources) 2956 1934 to 12,085

Sinks

Reaction with OH in troposphere –3180 j –2380 to –3970

Loss to stratosphere –200 a –100 to –300

Reaction with Cl in marine boundary layer –370 i, j –180 to –550

Microbial degradation in soil –180 k –100 to –1600

Loss to polar oceans –75 j, l, m –37 to –113

Subtotal (Sinks) –4005 –2797 to –6533

Total (Sources + Sinks) –1049 –4599 to +9288

The destructive flux from reaction with OH is calculated with revised OH (see Section 1.4).

a Khalil et al. (1999), adjusted for calibration.

b Lobert et al. (1999). c Yokouchi et al. (2002).

d Watling and Harper (1998). e Rhew et al. (2000).

f Varner et al. (1999a). g Dimmer et al. (2001).

h McCulloch et al. (1999); Keene et al. (1999).

i Redeker et al. (2000).

j Khalil and Rasmussen (1999), after calibration adjustment applied; see text.

k Khalil and Rasmussen (2000), after calibration adjustment applied; see text.

l Moore et al. (1996). m Unlike in Table 1-9, loss to polar oceans here is from field observations alone and is given as a net sink.
In addition to the methyl chloride source reported

by Yokouchi et al. (2002), Rhew et al. (2000) found large

net CH3Cl fluxes from halophytic plant species in the

middle and upper vegetation zones of Californian salt

marshes. Assuming that the sites were representative of

salt marshes globally, the authors postulated a global flux

of 170 Gg yr-1 from this source (Table 1-10). Subsequent

investigations on shrubland ecosystems in Southern

California demonstrated significant emissions from some

plant species at particular seasons (Rhew et al., 2001).

However, such emissions were counterbalanced by soil

uptake during other seasons, and therefore the biome was

in fact a small net sink for CH3Cl. Fluxes of CH3Cl from

peatland ecosystems have been monitored by investigators in Northeastern United States (Varner et al., 1999b)

and Western Ireland (Dimmer et al., 2001), although in

neither case was it clear whether emissions came from

plants or the soil. On the basis of these measurements

Varner et al. (1999b) postulated global emissions from

wetlands of 48 Gg yr-1, whereas Dimmer et al. (2001) estimated global emissions of 5 Gg yr-1 from peatlands and

35 Gg yr-1 from wetlands. These estimates must be treated

with caution because data in each investigation were

restricted to ecologically similar temperate sites and were

collected over a short period. Emissions of CH3Cl have

also been noted from the soil of California rice paddies,

which have been extrapolated to global emissions of 5 Gg

yr-1 from this crop (Redeker et al., 2000).

A novel abiotic mechanism for formation of CH3Cl

in soils rich in organic matter in the presence of Fe3+ and

high chloride concentrations has recently been proposed

(Keppler et al., 2000). Such a process may yield significant fluxes in highly saline environments, but no field

measurements have yet been conducted. Anthropogenic

release of CH3Cl during coal combustion, incineration,

and industrial processes was quantified in the Reactive

Chlorine Emissions Inventory (McCulloch et al., 1999).

Emissions from these sources contribute in all about 160

Gg yr-1

, or less than 10% of total known sources (Table

1-10).

1.5.2.3 SINKS OF ATMOSPHERIC METHYL CHLORIDE

Reaction with OH radical is the dominant pathway

for removal of CH3Cl from the atmosphere. Khalil and

Rasmussen (1999) estimated that the atmospheric lifetime

of CH3Cl with respect to OH attack varied from 0.8 years

in the tropics to 12 years at the poles, with an average

global lifetime of 1.4 years. This partial lifetime has been

revised to 1.5 years in this Assessment based upon the

considerations discussed in Section 1.4.2. Reaction with

chlorine radicals in the marine boundary layer also could

constitute another loss process, particularly in tropical latitudes (Keene et al., 1996). Using a chlorine concentration of 5 ¥ 10–4 radicals cm–3 in the marine boundary layer,

Khalil and Rasmussen (1999) estimated this sink as up to

0.4 Tg yr-1. This would represent a partial atmospheric

lifetime of 13 years with respect to this loss alone.

Some evidence exists for a potentially substantial

sink of atmospheric CH3Cl involving microbially mediated uptake by soil (Khalil and Rasmussen, 2000). On

the basis of measurements of CH3Cl uptake by soils from

Brazil and Greenland, these authors estimated a global

uptake of 0.18 Tg yr-1 (0.2 Tg before adjusting for calibration discrepancies) by soils. This number is uncertain,

and may be an underestimate; the ubiquitous occurrence

in soil from pristine environments of microbial species

capable of utilizing CH3Cl as sole carbon and energy

source implies a large and substantial soil sink for CH3Cl

(Harper, 2000; McAnulla et al., 2001). The loss of 0.2 Tg

yr-1 (0.18 Tg yr-1 adjusted for calibration) translates to a

partial atmospheric lifetime of 28 years.

Moore et al. (1996) noted undersaturations of

CH3Cl in the polar oceans, which Khalil and Rasmussen

(1999) combined with other data to calculate a global,

polar sink of 83 Gg yr-1 (75 Gg yr-1 adjusted for calibration). Uncertainties in the sink term given by Moore et

al. (1996) suggest that this value could range from 37 to

113 Gg yr-1. This small sink, if irreversible, amounts to

an atmospheric lifetime of ~70 years.

Combining the lifetimes of 1.5, 13, 28, and 70 years

for losses to reaction with OH, reaction with Cl, consumption in soils, and degradation in seawater, respectively,

yields an atmospheric lifetime of 1.3 years for CH3Cl

(Table 1-3).

1.5.2.4 ATMOSPHERIC BUDGET

The imbalance between calculated emissions from

known sources and the estimated global sink for atmospheric CH3Cl is about 1.0 Tg yr-1 out of a total estimated

flux of 4.0 Tg yr-1 (or 25%). All sources quantified to

date are situated predominantly in the tropics, a location

consistent with source deconvolution calculations. Using

a global 3-D model of atmospheric CH3Cl, Lee-Taylor et

al. (2001) were able to reproduce the observations of

Khalil and Rasmussen (1999) by adding a tropical terrestrial source of 2.5 Tg yr-1 and reducing emissions from

Southeast Asia. Any major as-yet-unidentified source

must necessarily show a similar global distribution to be

consistent with observations. In this context observations

of high CH3Cl mixing ratios in air masses associated with

forested tropical coastlines (Yokouchi et al., 2002) are

highly significant and should provide a focus for future

investigations on potential sources of atmospheric CH3Cl.

SOURCE GASES

1.43
1.6 ATMOSPHERIC HALOCARBON OBSERVATIONS COMPARED WITH EXPECTATIONS

1.6.1 Consistency Between Atmospheric

Halocarbon Measurements and

Known Sources and Sinks

Consistency between atmospheric measurements

and known sources and sinks for halocarbons are assessed

here in two complementary ways: (1) atmospheric measurements are compared with mole fractions that are calculated from emissions derived from industrial production data (see Section 1.3), and (2) the industry-based

annual emission estimates are compared with global emissions inferred from the atmospheric measurements. The

measurements we consider here are global means estimated from multiple surface sampling sites (the AGAGE,

CMDL, and UCI monitoring networks; D.R. Blake et al.,

1996; Montzka et al., 1999; Prinn et al., 2000) and global

means inferred from measurements at a single site (the

Australian air archive, Langenfelds et al., 1996; Fraser et

al., 1996). Calculations were performed with twodimensional models (Cunnold et al., 1994, 1997; Fraser

et al., 1999; Oram et al., 1995). Any unexpected differences between the calculated and observed mixing ratios

or the calculated and the industry-based emission estimates may represent unaccounted for production and

emissions, but these differences may also arise from systematic errors in measurement calibration or consistency,

in halocarbon loss rates (lifetimes), or in modeling the

atmosphere.

1.6.1.1 CFCS

Measured mixing ratios for CFC-11 and -12 are 5-

10% higher than those calculated from emissions derived

from production data throughout the entire period of

measurement (Figure 1-13). For CFC-12 during the 1990s

the discrepancy became small compared with the uncertainties in this analysis. For both these gases, mixing

ratios reported by different laboratories agree quite well

(Table 1-1), and this consistency argues against measurement error being the source of the discrepancies.

The differences noted for mixing ratios are also

apparent in a comparison of emissions derived from measurements or industry production data before 1992 (Figure

1-14). The observations before that year suggest approximately 5-10% larger emissions of CFC-11 and -12 compared with those derived from industry production data.

After 1992 the uncertainties in the analysis become large,

and any discrepancies are well within this uncertainty. In

the case of CFC-11, these differences also could be reconciled if its lifetime were somewhat longer than the current best estimate of 45 years (see Section 1.4). For CFC12, scaling the lifetime by some constant factor cannot

remove these differences; in fact it seems most likely that

the industry emissions estimates from 1993 to 1997 are

slightly high. This is consistent with improved containment of CFCs in applications (UNEP, 1998a), although

the extent to which practices to reduce ODSs emissions

have been adopted is difficult to ascertain.

Past discussion regarding CFC-113 noted that

observations substantially underestimated calculated

global means (Fraser et al., 1996; Prinn and Zander et al.,

1999; Figure 1-13). This situation stems from the presumption in these studies that significant amounts of CFC113 were produced by companies not reporting to AFEAS.

However, it is now apparent that the only significant

source that is not counted by AFEAS is Russian production (Bingfeng et al., 2000). The actual Russian production is unknown, but capacity for CFC-113 amounted to

some 12,000 tons yr-1 prior to the total shutdown that was

accomplished there in 2000 (personal communication, E.

Pedersen, World Bank, 2002). This represents a maximum of 5% of the 1989 reported production, with resulting emissions significantly less than the confidence limits

on global emissions (Figure 1-14). The global emissions

inferred from measurements are generally consistent with

AFEAS data accounting for global production of this gas.

There is a small divergence after 1995 for reasons not currently understood (Figure 1-14).

1.6.1.2 METHYL CHLOROFORM

For methyl chloroform artificially good agreement

is shown in Figure 1-13, because industry release estimates are used here to derive the OH field and methyl

chloroform lifetime (Prinn et al., 2001). An independently calculated OH field by Spivakovsky et al. (2000)

results in a methyl chloroform lifetime within 10% of that

inferred by Prinn et al. (2001), but a number of other

models of the OH distribution give much larger differences (e.g., Hein et al., 1997; Hauglustaine et al., 1998).

The small discrepancy between the measured and

expected mixing ratios after 1996 has been interpreted as

indicating lower OH values at that time (Prinn et al.,

2001). On the basis of an uncertainty of approximately

3% (except during 1996-1999) for industry-based emissions of methyl chloroform, it has been concluded that the

effective OH in a 12-box, 2-D model averaged from 1979

to 2000 has an uncertainty of ±14% (Prinn et al., 2001).

1.6.1.3 CARBON TETRACHLORIDE

With loss to the ocean now included (Yvon-Lewis

and Butler, 2002), the lifetime of carbon tetrachloride is

SOURCE GASES

1.44
SOURCE GASES

1.45

1975 1985 1995 2005

YEAR

100

200

300

400 MOLE FRACTION (ppt)

1975 1985 1995 2005

YEAR

200

400

600

800 MOLE FRACTION (ppt)

1975 1985 1995 2005

YEAR

50

100

150

200 MOLE FRACTION (ppt)

1975 1985 1995 2005

YEAR

50

100

150

200 MOLE FRACTION (ppt)

1975 1985 1995 2005

YEAR

50

100

150

200 MOLE FRACTION (ppt)

1975 1985 1995 2005

YEAR

50

100

150

200 MOLE FRACTION (ppt)

CFC-11 CFC-12 CFC-113

Methyl

chloroform

Carbon

tetrachloride

HCFC-22

Figure 1-13. Measured and calculated global mean surface mixing ratios for major CFCs, methyl chloroform,

carbon tetrachloride, and HCFC-22. Observations are from AGAGE in situ instruments (green lines: Prinn et

al., 2000), from CMDL in situ instruments and flask measurements (red lines: Montzka et al., 1999), and from

UCI flask measurements (purple lines: D.R. Blake et al., 1996, all gases except HCFC-22). For HCFC-22, the

AGAGE global means before 1992 are derived from measurements of the Australian air archive (Miller et al.,

1998) and a 12-box model calculation. Global surface mixing ratios were calculated with a 12-box model

(Cunnold et al., 1997) using global emissions from a combination of AFEAS (2001) and UNEP (2002) data and

lifetimes of 45 (29-76), 100 (77-185), 85 (54-143), 5.0 (4.5-5.6), and 12.0 (10.3-13.7) years for CFC-11, CFC12, CFC-113, methyl chloroform, and HCFC-22, respectively (black lines). Uncertainties in calculated mixing

ratios (shaded gray area) were derived from the lifetime ranges indicated, uncertainties in model parameters,

and uncertainties in industry emissions estimates. No calculated mixing ratios are shown for carbon tetrachloride because an independent record for historic production and emission of this gas is lacking. For CFC-113

mixing ratios have also been calculated (black dashed line) based on global emission estimates by Fisher et

al. (1994); corresponding calculations for the other gases are not shown because the results differ by less than

1% from the latest results.
SOURCE GASES

1975 1985 1995 2005

YEAR

0

200

400

600 SMOOTHED EMISSIONS (10 g/yr) SMOOTHED EMISSIONS (10 g/yr) SMOOTHED EMISSIONS (10 g/yr) SMOOTHED EMISSIONS (10 g/yr)

1975 1985 1995 2005

YEAR

0

200

400

600

1975 1985 1995 2005

YEAR

0

100

200

300

1975 1985 1995 2005

YEAR

0

100

200

300

CFC-11 CFC-12

CFC-113 Carbon tetrachloride

Figure 1-14. Estimates of annual emissions from industry (full black lines: based upon data from AFEAS

(2001) and UNEP (2002); see Section 1.3) and inferred from measurements (green lines and red symbols) for

CFC-11, CFC-12, CFC-113, and carbon tetrachloride. Uncertainties on industry emissions for CFC-11 are

from McCulloch et al. (2001); for CCl4, see UNEP (1998b) and McCulloch et al. (2002). The black dashed line

for CFC-113 indicates previous global emissions estimates that included an allotment from companies not

reporting to AFEAS (Fisher et al., 1994). The only available global emission estimate for CCl4 derived from

industry data was made by UNEP (1998b) for 1996 and is indicated by a black symbol. Release estimates

based on atmospheric measurements by AGAGE (green lines: Prinn et al., 2000) were made with a 12-box

model (Cunnold et al., 1997) and lifetimes of 45 (29-76), 100 (77-185), 85 (54-143), and 26 (17-36) years for

CFC-11, CFC-12, CFC-113, and carbon tetrachloride, respectively. The effect of these lifetime uncertainties,

which were derived from an ensemble of model calculations reported in WMO (1999), and, for CCl4, uncertainty in the oceanic loss (Yvon-Lewis and Butler, 2002), is indicated by the green error bars. Emission

estimates derived using the same AGAGE model applied to CMDL (Montzka et al., 1999) flask sample measurements (red diamonds) and CMDL in situ measurements (red asterisks) are also shown. All emission estimates derived from atmospheric measurements have been smoothed over 3-yr intervals to reduce the effects

of measurement imprecision. Units of emission are 109

g yr-1(109g = 1 Gg).

1.46
estimated to be 26 (17-36) years, or about 25% shorter

than in Prinn and Zander et al. (1999) (see Section 1.4.2).

Emissions estimated for 1996 by TEAP (UNEP, 1998b)

are below those inferred from measured trends and this

lifetime (Figure 1-14), and suggest that either (1) releases

of CCl4 are higher than estimated, or (2) the current, bestestimate lifetime is too short (Figure 1-14). Calibration

differences are too small to account for these large discrepancies (Table 1-1).

Uncertainties in lifetime estimates for CCl4 are significant and account for much of the uncertainty range

shown in Figure 1-14. Some insights into emissions of

CCl4 can be gleaned from a study of their relationship to

CFC production, which was believed to account for the

majority of carbon tetrachloride emissions (Simmonds et

al., 1998a). The molar ratio, [total global CCl4 emissions]/[total global production of CFC-11 and CFC-12],

ranges between 0.12 and 0.16 during 1978-1993 when the

global CCl4 emissions are inferred from atmospheric

measurements using a lifetime of 26 years (see Figure 1-

15). Some fraction of global CCl4 emissions, however,

arises from processes not related to CFC production;

TEAP (UNEP, 1998b) estimated that fraction at 30% in

1996. Such emissions, if included as a constant in other

years, suggest fugitive emissions of CCl4 arising from

CFC production of 11 ± 0.6% during 1978-1993. After

1993 as global CFC production declines, however, this

ratio increases dramatically (Figure 1-15). Such a change

suggests either that fugitive releases of CCl4 during CFC

production have increased by a factor of between 1.5 and

2 in recent years, or that CCl4 emissions not related to

CFC production are much higher than previously realized.

1.6.1.4 HCFCS

In the past there have been large discrepancies

between measured and calculated mixing ratios of HCFC142b and, to a lesser extent, HCFC-141b (Montzka et al.,

1994; Oram et al., 1995; Simmonds et al., 1998b). New

industry-based emission estimates (see Section 1.3;

AFEAS, 2001), however, result in much better agreement

for HCFC-142b (Figures 1-16 and 1-17). For HCFC-142b

and -141b, emissions inferred from measurements are still

slightly larger than the industry estimates, whereas for

HCFC-22 the industry releases appear to be on the high

side, and a somewhat shorter lifetime (more OH) would

result in a better fit. The incorporation of the timedependent OH field derived by Prinn et al. (2001) worsens

the differences between industry-derived emissions and

those calculated from atmospheric observations of HCFC22 (Figure 1-16). The effect of these potential OH changes

on the other HCFC plots is negligible, and is not shown.

1.6.1.5 HALONS

In the previous Assessment (WMO, 1999), substantial differences were noted between measured trends

and mixing ratios of Halon-1211 and those derived from

production data and release functions (Fraser et al., 1999;

Butler et al., 1998). Since then TEAP (an update to UNEP,

1999) has made global estimates of emissions for both

Halon-1211 and -1301 from independently derived functions that relate production to emission. For both halons,

mixing ratios calculated from these two emissions histories are substantially different but provide a better picture

of the uncertainties involved with this calculation than

were possible previously (Figure 1-17). Fairly significant

differences are also apparent in results from different laboratories for both halons: inter-laboratory differences

range up to 15% for Halon-1211 and 30% for Halon-1301.

SOURCE GASES

1.47 CCl Emission / CFC Production 4 Global CFC Production (ODP-kTon0.50 )

0.40

0.30

0.20

0.10

0.00

600

500

400

300

200

100

0

1970 1980 1990 2000

Figure 1-15. The molar ratio of global carbon tetrachloride emissions relative to aggregated CFC-11

and -12 global production (dashed line with open circles; left axis scale). In an attempt to better describe

the fraction of CCl4 emissions resulting from CFC

production, this ratio is calculated alternatively by

subtracting 22 Gg of CCl4 emissions for all years

(solid line with crosses; left axis scale) based upon

UNEP (1998b) estimates that 30% of CCl4 emissions

in 1996 were unrelated to CFC production in 1996

(UNEP, 1998b). In this analysis global emissions for

carbon tetrachloride were inferred from measurements (Prinn et al., 2000; Montzka et al., 1999) presuming a 26-yr global lifetime (Yvon-Lewis and

Butler, 2002; Section 1.4.2). Also shown is the global

aggregated CFC-11 and CFC-12 production in

recent years (thin black line, right axis scale; UNEP,

2002).
SOURCE GASES

1.48

1975 1985 1995 2005

YEAR

0

50

100

150 ESTIMATED EMISSIONS (10 g/yr)

HFC-134a

1975 1985 1995 2005

YEAR

0

20

40

60

80

100 ESTIMATED EMISSIONS (10 g/yr)

HCFC-141b

1975 1985 1995 2005

YEAR

0

10

20

30

40

50 ESTIMATED EMISSIONS (10 g/yr)

HCFC-142b

1975 1985 1995 2005

YEAR

0

100

200

300

400 ESTIMATED EMISSIONS (10 g/yr)

HCFC-22

Figure 1-16. Estimates of emissions from industry (solid black lines) and inferred from measurements (green

lines, red and blue symbols) for the most abundant HCFCs and HFC-134a. The green lines indicate global

release estimates from measurements at one to three sites and the AGAGE model (Cunnold et al., 1997; Prinn

et al., 2000; Miller et al., 1998) using global lifetimes of 14.0 years for HFC-134a, 9.3 years for HCFC-141b, 17.9

years for HCFC-142b, and 12.0 years for HCFC-22. The effect of an uncertainty of ±14% in the tropospheric

lifetime of these compounds due to hydroxyl (see Prinn et al., 2001) is indicated by the green error bars. The

red diamonds are annual release estimates obtained using the same AGAGE model applied to CMDL flask

measurements (7-8 sites; Montzka et al., 1999). The blue triangles are estimates obtained from an independent

set of measurements (at 41°S only) and a separate model (UEA: Oram et al., 1995, 1996). The black dashed

line on the HCFC-22 panel indicates emissions inferred from AGAGE measurements with the time-dependent

OH burden suggested by Prinn et al. (2001). Units of emission are 109 g yr-1 (109 g = 1 Gg).
SOURCE GASES

1.49

1975 1985 1995 2005

YEAR

0

2

4

6

HALON-1211

1975 1985 1995 2005

YEAR

0

1

2

3

4

HALON-1301

1975 1985 1995 2005

YEAR

0

5

10

15

20

HCFC-141b

1975 1985 1995 2005

YEAR

0

5

10

15

20

MOLE FRACTION (ppt)

HCFC-142b

MOLE FRACTION (ppt)

MOLE FRACTION (ppt)

MOLE FRACTION (ppt)

Figure 1-17. Observed and calculated global mean surface mixing ratios for selected halons and HCFCs over

time. Observations are global means from CMDL flask measurements (open red diamonds: Butler et al., 1998;

Montzka et al., 1999); global means from AGAGE in situ instruments (green crosses or solid green lines: Prinn

et al., 2000; Sturrock et al., 2001); global means from UCI quarterly flask measurements (open purple squares:

Blake et al., 2001, Halon-1211 only); and global means derived from an analysis of the Cape Grim Air Archive

by UEA (blue triangles: Fraser et al., 1999; Oram et al., 1995). Mixing ratios were calculated with 2-D models

and lifetimes of 17 years for Halon-1211 and 62 years for Halon-1301 (Fraser et al., 1999), 9.3 yrs for HCFC141b, and 17.9 years for HCFC-142b. Mixing ratios calculated with emissions of halons derived from industry

production data are from Fraser et al. (1999) (black dashed lines, halons only) and updated UNEP (1999) (gold

line). Mixing ratios of HCFCs are calculated from a combination of AFEAS (2001) and UNEP (2002) emissions

with updated release functions for HCFC-142b from foams (black lines; see Section 1.3), and, for comparison,

with the older release function for HCFC-142b from foams (dashed blue line, HCFC-142b panel only).
Within these rather large uncertainties there is broad consistency between calculated and observed mixing ratios

for both halons.

1.6.1.6 REGIONAL EMISSIONS

Initially using a simple long-range transport model

(Simmonds et al., 1996) and more recently a Lagrangian

dispersion model (Ryall et al., 2000), the European emission source strengths required to support the AGAGE

observations of pollution episodes at Mace Head, Ireland,

have been determined (see Table 1-11). Simmonds et al.

(1996) compared the pre-1995 estimates with industryderived emission estimates for Europe. They showed

agreement of the values within a factor of 2 and, in particular, roughly similar declines in the emissions from 1987

to 1994. These studies demonstrated the almost complete

phaseout of European emissions of methyl chloroform

and CFC-113. Although the releases of CFC-11 and -12

have dramatically declined, by more than a factor of 10,

European emissions of these gases are still continuing at

detectable and non-negligible levels. The contributions

of source regions to Mace Head fall off rapidly to the east

and south, with contributions falling by more than 3 orders

of magnitude toward the eastern Mediterranean, which is

a result of both fewer transport events to Mace Head and

greater dilution during transport to Mace Head (Ryall et

al., 2000).

1.6.2 Are Atmospheric Measurements

Consistent with Compliance with the

Fully Amended Montreal Protocol?

Atmospheric measurements can provide an indirect

and independent assessment of production and consumption data reported to UNEP. This assessment can then be

considered in light of the magnitude of known production

reported to UNEP and limits to production and consumption outlined in the Montreal Protocol. Significant uncertainties in this approach arise from uncertainties in

measurements and trace gas lifetimes, in model calculations of mixing ratios and emissions, and in the release

functions that relate emissions to industry production data.

Global measurements can provide estimates of

global emissions (Figures 1-14 and 1-16) and hence an

indication of the accuracy of global production and consumption data, but this analysis addresses compliance only

for ODSs that are restricted worldwide. Through 2001,

limits to production and consumption of ODSs in both

developed and developing countries applied only to CFCs.

By 2005, global limits to production and consumption by

Parties to the Protocol will apply also to halons, methyl

chloroform, carbon tetrachloride, and methyl bromide.

Global production of ODSs declined dramatically

during the 1990s because of limits set by the Protocol on

production in developed (non-Article 5(1)) countries

where most ODSs had occurred. In the data reported to

UNEP, production in non-Article 5(1) countries was

below the allowable limits throughout the 1990s (Figure

1-18).

As mentioned above, global limits apply only to

CFC production and consumption at the present time;

reported global production in 1999 was less than the

allowed consumption during that year by about 20 ¥ 103

ODP-tons. Measurements of all three major CFCs suggest 5-10% higher mixing ratios than expected based on

reported global production in 1999 (Figure 1-13). These

elevated mixing ratios do not necessarily imply additional

unreported production of these gases in 1999; for CFC-11

and -12, the mismatch is present throughout the entire

measurement period and unaccounted-for emissions in

recent years cannot explain this discrepancy. In fact,

observed and calculated mixing ratios for CFC-12 have

been converging during the 1990s, indicating that emissions estimated by industry since 1995 may be overestimates by an average of about 30 ¥ 103 ODP-tons yr-1 (or

30 Gg; Figure 1-14), consistent with improved containment in recent years (UNEP, 1998a). For CFC-113 good

consistency is observed until the beginning of the 1990s.

After that a small divergence becomes apparent. This mismatch could be reconciled with about 10 ¥ 103 ODP-tons

of additional CFC-113 emission in recent years. If the

small discrepancies for CFCs noted here were all attributed to errors in reported production, the measured atmospheric trends would suggest slightly less aggregate CFC

production in the late 1990s than reported to UNEP. CFC

production reported to UNEP in 1999 was somewhat (~20

¥ 103 ODP-tons) below the Montreal Protocol limit for

that year (UNEP, 2002); the atmospheric measurements

are not inconsistent with this assertion.

For halons, production data reported to the UNEP

Ozone Secretariat for 1995-1999 are similar to the global

limit that will apply beginning in 2002. If the trends indicated by the data in 1998 and 1999 persist into the future,

production would be quite below the limits allowed in

2002 (Figure 1-18). Assessing the accuracy of reported

production with atmospheric measurements for halons

(Figure 1-17) is difficult because of the rather large uncertainties in measurement calibration, in the relation

between production and emission for halons, and in the

atmospheric lifetime for Halon-1211. Within these rather

large uncertainties the atmospheric observations are

broadly consistent with known production reported to

UNEP for both Halon-1211 and Halon-1301.

SOURCE GASES

1.50
For carbon tetrachloride, difficulties have arisen in

reporting appropriate data to UNEP (UNEP, 1998b). In

light of this, limits to production and consumption are

compared with emissions derived from measurements

rather than consumption or production data reported to

the UNEP Ozone Secretariat (Figure 1-18). For most uses,

emissions of CCl4 closely follow sales, so this approach

can yield some insight. Global limits on CCl4 production

apply only after 2005, so it is not meaningful to ask if

atmospheric measurements suggest compliance until after

that date. Inferred emissions in 1999, however, are about

7 times greater than the limits to global production set for

2005.

For methyl chloroform, global production reported

in 1999 was substantially lower than the limit for nonArticle 5(1) countries. In addition, production in 1999

was already comparable with the global production limit

for 2002.

For HCFCs, limits to global production apply only

after 2015. Global production is below the current consumption limit for non-Article 5(1) countries (Figure 1-18).

1.7 OTHER TRACE GASES

1.7.1 Carbon Dioxide (CO2)

Changes in climate and meteorological conditions

affect the ozone layer, because ozone depletion and climate

change share common physical and chemical processes.

Anthropogenic carbon dioxide (CO2) is the major contributor to the increase in radiative forcing from trace gas

changes since preindustrial times. For this reason, atmospheric CO2, the global carbon cycle, and mankind’s impacts

on them through such activities as fossil fuel combustion

and changes in land use are addressed in considerable detail

in Chapter 3 of the IPCC Climate Change 2001 report

(Prentice et al., 2001). Similarly, the 2001 IPCC Special

Report on Emissions Scenarios, or SRES (Nakic ́enovic ́ et

al., 2000) presents a wide range of projected emissions of

CO2 and other radiatively important trace gases over the

coming century. We present here only a very brief overview of the main points, and we refer the reader to those

IPCC reports for detailed discussions and references.

SOURCE GASES

1.51

Table 1-11. European emission source strengths required to support the AGAGE observations of each

halocarbon and greenhouse gas at Mace Head, Ireland, over the period 1987-2000 (emissions in thousand tons per year).

Year CFC-11 CFC-12 CFC-113 CH3CCl3 CCl4 CHCl3 CH4 N2O

1987 215 153 68 207 27 47 900 2 200

1988 132 98 43 115 20 22 500 1 800

1989 94 87 63 159 26 22 700 1 300

1990 67 64 63 153 12 31 500 1 800

1991 52 59 44 108 14 35 200 2 700

1992 27 32 31 88 9 31 000 1 600

1993 20 22 21 75 6 23 600 1 600

1994 25 31 20 115 5 40 900 2 400

1995 9 15 5 31 3 20 24 800 1 400

1996 8 14 4 19 4 15 27 800 1 500

1997 9 13 2 8 3 19 27 300 1 700

1998 10 13 2 2 3 20 25 800 1 600

1999 9 11 1 0 2 19 32 000 1 600

2000 6 6 1 0 1 16 20 500 1 000

HCFC-141b HCFC-142b HCFC-22 HFC-134a HFC-152a

1995 8 6 23 3 0.4

1996 9 6 20 6 0.5

1997 13 9 34 11 0.7

1998 14 8

1999 12 8 29 13 1

2000 9 4 20 10 0.7

A simple long-range transport model (Simmonds et al., 1996) was employed for the 1987-1994 period and a Lagrangian dispersion model (Ryall et al.,

2000) thereafter.
SOURCE GASES

1.52

0

200

400

600

800

1000

1200

1970 1980 1990 2000 2010

0

20

40

60

80

100

120

140

160

1970 1980 1990 2000 2010

0

10

20

30

40

50

60

70

80

1970 1980 1990 2000 2010

0

5

10

15

20

25

30

35

40

45

1970 1980 1990 2000 2010

0

20

40

60

80

100

120

140

160

180

200

1970 1980 1990 2000 2010

0

5

10

15

20

25

30

35

40

45

50

1970 1980 1990 2000 2010 CFC Production (ODP-kTon) CCl4 Inferred Emission vs Consumption Limits (ODP-kTon) CH3CCl3 Production (ODP-kTon) Halon Production (ODP-kTon) CH3Br Production (ODP-kTon) HCFC Cons. or Prod. (ODP-kTon)

Figure 1-18. A comparison between reported global ODS production (filled solid diamonds with connecting

lines: AFEAS, 2001; UNEP, 2002); ODS production reported in non-Article 5(1) countries (open circles with

connecting lines: UNEP, 2002); and global ODS production limits (thick solid lines) as outlined in the fully

revised and amended Montreal Protocol. Also shown are production or consumption limits for non-Article 5(1)

countries in years before global limits apply (short-dashed lines). For CCl4, global emissions inferred from

measurements are shown instead of reported global production. For HCFCs, global limits to production and

consumption apply only after 2015, so the non-Article 5(1) consumption limit (short-dashed lines) and production limit (thick long-dashed line) are shown instead. For some ODSs, global production in the late 1990s is

higher than non-Article 5(1) limits because significant amounts were produced in countries operating under

Article 5(1) during those years.
The atmospheric burden of CO2 has increased by

about 35% since 1750, to a mean abundance of ~370 parts

per million (ppm) in 2000 (Table 1-12). Systematic measurements of atmospheric CO2 since the late 1950s show

a regular increase, with a mean growth rate at Earth’s

surface of ~1.5 ppm yr-1 (0.4% yr-1) over the past two

decades (Figure 1-19; Tans et al., 2002). This increase is

roughly half of that expected from the combustion of fossil

fuels and cement production, the difference being due to

CO2 uptake by the oceans and by the terrestrial biosphere.

There is considerable interannual variability in this rate

of increase that is due to the effects of climate variability

(e.g., El Niño events) on these uptake processes (Battle et

al., 2000; Le Quere et al., 2002). This variability is also

shown by measurements of total CO2 column abundances

above the Jungfraujoch (Zander et al., 2002).

CO2 from fossil fuel burning will continue to be

the dominant factor determining atmospheric CO2 levels

during the 21st century. Among the six emissions scenarios considered by the SRES (Nakic ́enovic ́ et al., 2000),

the projected range of CO2 abundances in 2100 was

between 540 and 970 ppm, substantially higher than

present-day amounts (Figure 1-19). Among the largest

uncertainties in these projections are the effects of climate

change on rates of deep ocean mixing, as well as on surface ocean and terrestrial biological productivity.

1.7.2 Methane (CH4)

Atmospheric methane (CH4) is a greenhouse gas

whose radiative properties and atmospheric chemistry

affect both climate and stratospheric ozone. Because

methane is a sink for reactive chlorine and a source of

H2O vapor, it directly influences the availability of inorganic halogen for depleting stratospheric ozone. Furthermore, the atmospheric chemistry of methane affects

hydroxyl radical (OH) abundance on global scales. This

can alter degradation rates (lifetimes) of methane and

ozone-depleting substances such as methyl chloroform

and HCFCs. Finally, methane affects climate. Hansen

and Sato (2001) estimate a change in radiative forcing for

CH4 in 2000 compared with preindustrial time, including

direct and indirect effects, of 0.7 ± 0.2 W m-2, which is

one-half the change in forcing due to CO2 increases.

Although global concentrations of methane have

varied in preindustrial times, amounts present in today’s

atmosphere are thought to be higher than at any time in

the past 420,000 years. Whereas the globally averaged

methane mole fraction in 2000 was estimated to be

between 1752 ppb (updates of Dlugokencky et al., 2001)

and 1784 ppb (AGAGE network, Tohoku University

gravimetric scale; Cunnold et al., 2002) (Table 1-12),

measurements of CH4 in air extracted from polar ice indicate that its abundance during the last 420,000 years of

the preindustrial era varied between about 350 ppb during

glacial times and about 700 ppb during interglacial

periods (Petit et al., 1999). From ice-core measurements

in Antarctica and Greenland, and Antarctic firn and

archived air from southern midlatitudes, Etheridge et al.

(1998) showed that the globally averaged methane mole

fraction was relatively constant at 695 ppb from 1000 to

1800 A.D., and increased through the industrial era.

Systematic measurements of atmospheric CH4 show

that its global abundance at Earth’s surface continues to

increase (Figure 1-19); the average growth rate was 7.9 ±

0.1 ppb yr-1 from 1984 to 2000, and the rate of increase

slowed from about 14 ppb yr-1 in the early 1980s to less than

5 ppb yr-1 in the late-1990s (except during 1998). The 1999-

2000 global growth rate was less than 2 ppb yr-1 (Table 1-

12). Significant interannual variability is superimposed on

the long-term decrease in growth rate (Dlugokencky et al.,

1998, 2001; Simpson et al., 2002; Cunnold et al., 2002).

Monitoring of the vertical column abundance of CH4 above

the Jungfraujoch since the mid-1980s also confirms the continued slowing of the rate of CH4 accumulation in the atmosphere, which was found equal to 0.72% in 1987-1988,

0.51% in 1991-1992, 0.32% in 1995-1996, and 0.14% in

1999-2000 (Zander et al., 2002).

The cause of the decrease in methane growth rate is

not fully understood. Although the major sources of atmospheric methane are known, they are poorly quantified.

Bottom-up assessments from individual sources of emission rates and their changes over time are not reliable

enough to discern small changes in global emission rates.

Changes in hydroxyl radical abundance also would affect

methane growth rates, but much uncertainty remains

regarding trends of this important atmospheric oxidant

(see Section 1.4; Krol et al., 1998; Prinn et al., 2001).

Dlugokencky et al. (1998) have suggested that the recent

methane trends can be explained by an approach of the

system to steady state, i.e., fairly constant emissions and

lifetime in recent years. Resolution of these issues has

important implications for projecting future concentrations.

IPCC (Prather and Ehhalt et al., 2001) calculated

future CH4 abundances from 2000 to 2100 based on CH4

emission rates in the six illustrative scenarios in the IPCC

Special Report on Emissions Scenarios (Nakic ́enovic ́ et

al., 2000; Figure 1-19). By 2100, CH4 mixing ratios range

from about 1575 ppb (i.e., 10% below present-day mixing

ratios; 1880 ppb in 2050) for IPCC scenario B1, which

emphasized relatively small global population and

resource-efficient technologies, to 3730 ppb (2560 ppb in

2050) for scenario E0 (not shown in Figure 1-19), which

SOURCE GASES

1.53
SOURCE GASES

1.54

Table 1-12. Mixing ratios and growth rates for HFCs, FCs, sulfur hexafluoride, nitrous oxide, methane,

carbon dioxide, and carbon monoxide.

Species Common or Mixing Ratio Growth (1999-2000) Laboratory (No. of sites),

Industrial Name 1996 1998 2000 ppx yr–1 % yr–1 Method

HFCs (ppt)

CHF3 HFC-23 11.7 13.5 15.5 0.9 5.8 UEA, SH (1), flask a

CHF2CF3 HFC-125 0.92 1.4 0.3 22 AGAGE (2), in situ a

CH2FCF3 HFC-134a 7.9 14.6 3.5 27 AGAGE (2), in situ a

3.0 7.6 13.8 3.2 26 CMDL (8), flask a

CH3CHF2 HFC-152a 1.3 1.7 0.2 12 AGAGE (2), in situ a

FCs (ppt)

CF4 FC-14 76.0 MPAE, NH (1), flask a

CF3CF3 FC-116 2.2 UEA, SH (1), flask a

c-C4F8 FC-318 1.0 UEA, SH (1), flask a

SF6 (ppt) Sulfur hexafluoride 3.5 3.9 UH, SH (1), flask b

3.6 UEA, SH (1), flask a

3.8 4.3 4.7 0.21 4.6 CMDL (7), flask b

4.2 4.6 0.18 4.0 CMDL (5), in situ b

N2O (ppb) Nitrous oxide 312.7 314.1 316.0 1.0 0.32 AGAGE (4-5), in situ b

312.0 313.8 315.7 0.85 0.27 CSIRO, SH (1), flask b

312.2 313.2 315.2 1.0 0.32 CMDL (7), flask b

312.7 314.3 317.2 1.1 0.35 CMDL (5), in situ b

CH4 (ppb) Methane 1761.7 1774.7 1784.4 1.3 0.07 AGAGE (4-5), in situ c

1681.3 1695.8 1707.6 1.7 0.10 CSIRO, SH (1), flask c

1729.7 1743.2 1751.6 0.7 0.04 CMDL (42), flask c

1751.2 1764.7 1771.1 2.1 0.12 UCI (~40), flask c

CO2 (ppm) Carbon dioxide 361.6 365.4 368.6 1.1 0.31 CMDL (44), flask d, *

CO (ppb) Carbon monoxide 112.1 144.5 112.2 –11.6 –9.9 AGAGE, NH (2-3), in situ e

110.8 132.1 106.0 –13.5 –12.3 CMDL, NH (24), flask e

52.3 56.4 51.0 –2.3 –4.4 AGAGE, SH (2), in situ e

50.8 56.2 50.6 –2.2 –4.2 CSIRO, SH (1), flask e

54.3 60.3 55.9 –1.8 –3.2 CMDL, SH (19), flask e

Global mixing ratios (dry air mole fractions) and growth rates are at Earth’s surface unless otherwise specified as being from Northern Hemisphere

(NH) or Southern Hemisphere (SH) sites only. Laboratories: Advanced Global Atmospheric Gases Experiment (AGAGE); Global Atmospheric

Sampling Laboratory, Commonwealth Scientific and Industrial Research Organization (CSIRO), sampling at 41°S; Max-Planck-Institut für

Aeronomie (MPAE), sampling at Lindau, Germany; Climate Monitoring and Diagnostics Laboratory (CMDL), National Oceanic and Atmospheric

Administration (NOAA); University of East Anglia (UEA), sampling at 41°S; University of Heidleberg (UH), sampling at 41°S; University of

California at Irvine (UCI), quarterly sampling at about 40 sites.

Data sources: AGAGE: Prinn et al. (2000); Cunnold et al. (2002); ftp://cdiac.esd.ornl.gov/pub/ale_gage_Agage. CMDL: Dlugokencky et al. (2001);

Montzka et al. (1996b); Novelli et al. (1998); Tans et al. (2002); www.cmdl.noaa.gov. CSIRO: Langenfelds et al. (2002). MPAE: Harnisch et al.

(1999). UEA: Oram et al. (1998); Oram (1999). UCI: Simpson et al. (2002). UH: Maiss and Brenninkmeijer (1998).

* Many other groups measure CO2; only results from the most extensive network are included here (see Prentice et al., 2001). a Measurements by gas chromatography-mass spectrometric detection.

b Measurements by gas chromatography-electron capture detection. c Measurements by gas chromatography-flame ionization detection.

d Measurements by nondispersive infrared detection. e Measurements by gas chromatography-mercuric oxide reduction detection.
included rapid population growth through the century and

regional solutions to environmental problems. Though

the range of values is large, the scenarios provide a range

of potential future atmospheric composition for given

amounts of CH4 emissions.

1.7.3 Nitrous Oxide (N2O)

Nitrous oxide (N2O) is the major source of ozonedepleting nitrogen oxides (NOx) in the stratosphere. These

reaction products of N2O photochemistry make a significant contribution to both anthropogenic and natural ozone

depletion. New rate constants for some stratospheric reactions indicate an enhanced role for NOx in ozone loss

(Portmann et al., 1999). Consequently, it is possible that

enhanced future N2O emissions could noticeably delay

stratospheric ozone recovery (Randeniya et al., 2002).

The atmospheric burden of N2O has been

increasing steadily for decades (Figure 1-19). NOAA/-

CMDL and AGAGE data show that the globally averaged

SOURCE GASES

1.55

310

330

350

370

ppb

1980 2000 2020 2040

year

B 2

B1

A2 A1FI

700

900

2100

2300

2500

ppb

1980 2000 2020 2040

year

B 2

B 1

A 2

A1FI

370

410

450

490

530

ppm

1980 2000 2020 2040

year

A1FI

B2

B 1

A 2

c a bon dioxide ethane

nitrous oxide

m

Figure 1-19. Long-term trends (annual mean data)

and possible future levels of carbon dioxide (CO2),

methane (CH4), and nitrous oxide (N2O). Data are

global means from CMDL (red lines: Tans et al., 2001;

Dlugokencky et al., 2001; Hall et al., 2002), AGAGE

(green lines: Prinn et al., 2000; Cunnold et al., 2002),

and UCI (purple line: Simpson et al., 2002, methane

only); CSIRO Cape Grim data are from 41∞S only (light

blue lines: Langenfelds et al., 2002). The scenarios

(plotted from 2000) are from IPCC (2001; Appendix

II): A1F1, a future with rapid economic growth and

intensive use of fossil fuels; A2, slower economic

development, but with continuously increasing global

population; B1, progress toward economic and environmental sustainability, and clean, resource-efficient

technologies abound; B2, intermediate levels of economic development, continuously increasing global

population, and less rapid technological changes.
mixing ratio in January 2001 was 315-317 ppb (Prinn et

al., 2000; Hall et al., 2002). Its rate of increase has averaged ~0.75 ppb yr-1

, with some interannual variability

since the late 1970s. Its lifetime is currently estimated at

120 (97-137) years, although its “perturbation” lifetime is

4-6% shorter (Prather, 1998).

Most updates in our understanding of the sources

and sinks of atmospheric N2O since the 1998 Assessment

(WMO, 1999) were assessed in Chapter 4 of the Third

Assessment Report (TAR) for the IPCC (Prather and

Ehhalt et al., 2001). A number of studies discussed in that

report show a preindustrial tropospheric abundance of N2O

of ~270 ppb, suggesting an increase in the atmospheric

burden of about 17% since then. Little has changed in our

overall understanding of the atmospheric budget of this

gas since the 1998 Assessment. Known sources of 15-18

Tg yr-1, mainly from natural soils, agricultural activities,

biomass burning, and the ocean, can be accounted for by

known sinks and observed increases. However, the uncertainties in the source estimates are large. The suggested

need for a large stratospheric or exotic source of N2O to

account for the isotopic composition of the troposphere

has been found by a number of studies not to be necessary

(Prather and Ehhalt et al., 2001).

1.7.4 Carbon Monoxide (CO)

In the troposphere, carbon monoxide (CO) is

closely linked to the cycles of ozone and methane, and

plays an important role in the control of the OH radical

concentration. Hydroxyl radical, in turn, controls the lifetime of many reduced compounds (e.g., CH4, HCFCs,

HFCs, and CH3CCl3). Therefore, predicting future concentrations of these ozone-relevant and radiatively important gases depends partly on how well the CO budget and

its potential for change is understood. CO could be

assigned a GWP because of its effect on the lifetime of

other greenhouse gases (e.g., Prather, 1996; Daniel and

Solomon, 1998; Prather and Ehhalt et al., 2001). In this

regard, emitting 100 Tg of CO is equivalent to emitting 5

Tg of CH4 (Wild and Prather, 2000; Derwent et al., 2001).

Most budgets evaluated by Prather and Ehhalt et al. (2001)

estimate a current annual source of CO to the troposphere

in the range of 2000 to 3000 Tg.

1.7.4.1 ATMOSPHERIC GLOBAL DISTRIBUTION AND

TRENDS

Modern direct measurements indicate that trends

in atmospheric CO were positive until the mid- to late1980s at least in the Northern Hemisphere, but since then

the mean trend has been negative (Zander et al., 2000;

Khalil and Rasmussen, 1994; Mahieu et al., 1997; Novelli

et al., 1998), e.g., reaching a decrease of –0.6% yr-1 above

the Jungfraujoch in 1999-2000 (Zander et al., 2002). CO

exhibits both short- and long-term changes on regional

and global scales and, after an overall increase over several decades, it appears that CO has declined in many parts

of the world (Table 1-12; see also Kurylo and Rodríguez

et al., 1999). Causes of these recent changes have not

been identified with certainty, but decreases in tropical

biomass burning (Yung et al., 1999), decreases in Northern

Hemispheric urban emissions (Bradley et al., 1999), and

increases in OH concentrations due to stratospheric ozone

depletion (Krol et al., 1998; Law, 1999) may have played

a role (although one study suggests decreases in OH in

the 1990s (Prinn et al., 2000)). Short-term increases in

CO column abundances and surface mixing ratios were

observed during the strong El Niño of 1998, probably as a

result of enhanced burning then (Wotawa et al., 2001;

Rinsland et al., 2000). No long-term change has been

detected in CO mixing ratios at Cape Point, South Africa,

during 1978-1998 (Scheel et al., 1995).

1.7.4.2 CO SOURCES AND SINKS

CO has significant atmospheric photochemical

sources from the oxidation of CH4 and non-methane

hydrocarbons, as well as emissions from fuel combustion

and biomass burning. Considerable uncertainties exist in

the CO budget. Recent evaluations of global sources and

budget suggest that the principal sources are biomass

burning, fossil fuel combustion, agricultural waste

burning, biofuel combustion, and industrial processes.

The global surface source is estimated at 1550 Tg yr-1

(Prather and Ehhalt et al., 2001).

The most important sink of CO is reaction with OH.

Uncertainties in OH distributions and the reaction rate

constant of CO + OH result in a large range of estimates

for this sink. McCabe et al. (2001) have studied the

kinetics of CO + OH for atmospheric conditions in the

laboratory, decreasing the uncertainty in the rate coefficient significantly. Atmospheric CO is consumed by soils;

recent studies (Potter et al., 1996; Sanhueza et al., 1998;

King, 2000) suggest global consumption by soils of 100-

300 Tg yr-1

, lower than the 190-580 Tg yr-1 proposed by

Conrad and Seiler (1985), which has been widely used in

previous Assessment budgets.

1.7.4.3 ISOTOPIC DATA

Isotopic studies of atmospheric CO (Brenninkmeijer et al., 1999) provide an absolute technique by

which atmospheric mixing ratios of CO are determined

(Brenninkmeijer et al., 2001). Confirmation that CO

undergoes mass-independent fractionation (Röckmann et

SOURCE GASES

1.56
al., 2001) indicates that previous 13C data may need a

small, but significant, correction (Röckmann and

Brenninkmeijer, 1998). Additional work is needed in

defining fractionation of some CO production processes,

notably, CH4 and non-CH4 hydrocarbon oxidation, and its

combustion sources.

1.7.5 Carbonyl Sulfide (COS)

Carbonyl sulfide (COS) is the most abundant sulfurcontaining trace gas in the atmosphere. Because it has a

lifetime that is long relative to atmospheric mixing times

(>1 year), much of the COS emitted or produced in the troposphere reaches the stratosphere. Some fraction of stratospheric COS becomes photolyzed and contributes sulfur

to the stratospheric aerosol layer. This aerosol provides

surfaces that enhance the concentration of ozone-depleting

radicals; it also plays an important role in the radiative balance of the atmosphere. Earlier work had suggested that

COS might be the main source of nonvolcanic sulfate to

the stratosphere (Crutzen, 1976). More recent work, however, based on revised reaction rates, cross sections, and

photolytic quantum yields, has suggested that COS may

account for a smaller fraction of the sulfate present in the

stratospheric Junge layer than previously thought (Chin

and Davis, 1995; Kjellstrom, 1998; see also Chapter 2).

A combination of results from total column

abundances (Rinsland et al., 2002a; Mahieu et al., 1997;

Griffith et al., 1998), periodic surface measurements

(Bandy et al., 1992; Thornton et al., 1996), and an analysis

of firn air in both hemispheres (Sturges et al., 2001b) suggests that the atmospheric burden of COS was fairly constant from the mid-1900s to the early 1990s. In the 1990s

updated results suggest a slight decrease (about 1% yr-1

or smaller) (Mahieu et al., 1997; Sturges et al., 2001b);

Rinsland et al. (2002a) have inferred a mean tropospheric

decrease from 1978 to 2002 at 32°N equal to 0.25 ± 0.04%

yr-1

. Firn air results suggest that the recent decreases may

have been more pronounced in the Northern Hemisphere

than in the Southern Hemisphere (Sturges et al., 2001b);

they also point to the possibility of lower COS mixing

ratios in the oldest air sampled in Antarctica. Results from

an Antarctic ice core suggest that preindustrial mixing

ratios of COS were 25% lower than observed in the

modern atmosphere (Aydin et al., 2002).

Although numerous sources and sinks of COS have

been identified, much uncertainty remains regarding the

magnitude of these fluxes. Until recently, best estimates

of total sources outweighed sinks for COS by ~30%,

despite evidence to suggests fairly constant mixing ratios

over time in the atmosphere. More recent work by Kuhn

et al. (1999), Kesselmeier et al. (1999), and Simmons et

al. (1999) confirmed the earlier work of Castro and

Galloway (1991) and suggests a much enhanced role for

soils in the removal of COS from the atmosphere. With

this new understanding of soil losses and some revision in

the net influence of open and coastal oceans, Watts (2000)

suggested that the magnitude of known sinks is similar to

known sources despite the large uncertainties. The magnitude of sinks or sources and a mean atmospheric concentration of about 500 ppt imply a COS global lifetime of

about 4 years (Chin and Davis, 1995; Watts, 2000).

1.7.6 Hydrofluorocarbons (HFCs)

The most abundant hydrofluorocarbons (HFCs) in

the present-day atmosphere are HFC-23 (CHF3) and HFC134a (CH2FCF3). Since the previous Assessment, the tropospheric abundance of each has continued to increase.

In addition, three more HFCs have been identified in the

background atmosphere: HFC-125 (CHF2CF3), HFC143a (CH3CF3), and HFC-152a (CH3CHF2). Their atmospheric mixing ratios are relatively low but are increasing

at significant rates.

1.7.6.1 HFC-134a

HFC-134a (CH2FCF3) is used extensively as a

refrigerant and air-conditioning agent, and to a lesser

degree as a foam-blowing agent. Updated measurements

show that the global mean surface mixing ratio of HFC134a had risen to just under 15 ppt by mid-2000 (Table 1-

12; Figure 1-20; Simmonds et al., 1998b; Montzka et al.,

1999; Prinn et al., 2000; Sturrock et al., 2001). The current growth rate of about 3.0 ppt yr-1 has approximately

doubled since 1996. Calibration differences between

CMDL and AGAGE are less than 5%.

IPCC (Prather and Ehhalt et al., 2001) calculated

future HFC-134a abundances from 2000 to 2100 based

on HFC-134a emission rates in the six illustrative scenarios in the IPCC Special Report on Emissions Scenarios

(Nakic ́enovic ́ et al., 2000; Figure 1-20). By 2100, HCFC134a mixing ratios range from about 380 ppt (290 ppt in

2050) in IPCC scenario B1 to 960 ppt (535 ppt in 2050)

in IPCC scenario A1.

1.7.6.2 HFC-23

Although HFC-23 (CHF3) has found minor uses in

refrigeration, plasma etching, and as a fire extinguishant,

its major atmospheric source is as a byproduct from the

manufacture of HCFC-22 (Oram et al., 1998). In addition, Engen et al. (1998) reported elevated levels of HFC23 when sampling the emissions from an aluminum

production plant.

SOURCE GASES

1.57
The only available time series of tropospheric

HFC-23 measurements are those of Oram et al. (1998).

The mixing ratio of HFC-23 at Cape Grim has risen from

2 ppt in 1978 to 15.5 ppt in 2000 (Table 1-12; Figure 1-

20). Despite the emission reductions that have occurred

in developed countries as a result of improved HCFC-22

process control and capture of the HFC-23 byproduct, the

growth rate of HFC-23 at Cape Grim has increased in

recent years from 0.55 ppt yr-1 in 1995 to 0.9 ppt yr-1 (6%

yr-1) in late 1999. This may have resulted from increased

HCFC-22 production in developing countries recently

(McFarland and van Gerwen, 2000).

The presence of HFC-23 in the atmosphere has

been confirmed by Engen et al. (1998) and Culbertson et

al. (2000). In the latter study, a mean mixing ratio of 18 ±

2 ppt was measured in air samples collected at a rural,

continental site in the United States in March-April 2000.

IPCC (Prather and Ehhalt et al., 2001) calculated

future HFC-23 abundances from 2000 to 2100 based on

HFC-23 emission rates in the six illustrative scenarios in

SOURCE GASES

1.58

5

1 0

1 5

ppt

1980 2000 2020 2040

year

IS92a

50

100

150

ppt

1980 2000 2020 2040

year

B 2

B 1

A2 A 1

5

1 0

1 5

2 0

2 5

ppt

1980 2000 2020 2040

year

B2

B 1

A2

A1

10

20

30

ppt

1980 2000 2020 2040

year

A1,A2,B1,B2

HFC-23

HFC-134a

HFC-125

HFC-152a

Figure 1-20. Long-term trends (annual mean data) and possible future levels of HFCs. Data are global means

from CMDL (red line: Montzka et al., 1996b, HFC-134a only), and AGAGE (green lines: Prinn et al., 2000;

Sturrock et al., 2001). Cape Grim results are for 41∞S only (blue lines: Oram et al., 1996; 1998). The scenarios

(plotted after 2000) are from IPCC (2001; Appendix II).
the IPCC Special Report on Emissions Scenarios

(Nakic ́enovic ́ et al., 2000; Figure 1-20). By 2100, HCFC23 mixing ratios are about 32-34 ppt (35 ppt in 2050) in

all IPCC scenarios.

1.7.6.3 HFC-125, HFC-143a, AND HFC-152a

The major use of HFCs -125 (CHF2CF3) and -143a

(CH3CF3) is as refrigerants, whereas HFC-152a

(CH3CHF2) is used primarily as a foam-blowing agent and

aerosol propellant. The current tropospheric background

abundance of these HFCs is relatively low, in the range of

1-3 ppt (Table 1-12; Prinn et al., 2000; Culbertson et al.,

2000; Sturrock et al., 2001). Southern Hemispheric

growth rates in mid-2000, based on measurements at Cape

Grim, were about 0.2 ppt yr-1 and 0.1 ppt yr-1 for HFC125 and HFC-152a, respectively (Sturrock et al., 2001).

IPCC (Prather and Ehhalt et al., 2001) calculated

future HFC-125, -143a, and -152a abundances from 2000

to 2100 based on respective emission rates in the six illustrative scenarios in the IPCC Special Report on Emissions

Scenarios (Nakic ́enovic ́ et al., 2000; Figure 1-20; HFC143a not shown). By 2100, mixing ratios are about 60-

140 ppt for HFC-125 (30-60 ppt in 2050), 90-200 ppt for

HFC-143a (40-70 ppt in 2050), and 80 ppt for HFC-152a,

(55 ppt in 2050) for all these IPCC scenarios.

1.7.7 Fluorocarbons (FCs), SF6, and SF5CF3

Due to a combination of long atmospheric lifetimes

(>1000 years) and strong infrared absorption cross sections, fluorocarbons (FCs) are among the most potent

greenhouse gases known, having GWPs about 4 orders of

magnitude higher than that of CO2 (see Table 1-6). These

gases are useful tracers of atmospheric transport in both

the troposphere and the stratosphere. The only important

sinks for FCs and sulfur hexafluoride (SF6) are photolysis

or ion reactions in the mesosphere.

1.7.7.1 CF4 AND C2F6

The most abundant FCs in the present-day atmosphere are CF4 (FC-14) and C2F6 (FC-116). Approximately

50% of the current atmospheric burden of CF4 (~80 ppt)

was present in preindustrial times, implying a natural,

probably geochemical, source (Harnisch et al., 1996a;

Harnisch and Eisenhauer, 1998). The predominant source

of both compounds in recent times is believed to be the

aluminum industry. Other minor sources include the semiconductor industry and from the production of CClF2CF3

(CFC-115), where C2F6 can be formed as a byproduct due

to overfluorination.

The current background atmospheric mixing ratio

of CF4 is about 80 ppt (Table 1-12; Harnisch et al., 1999;

Culbertson et al., 2000). The measurements of Harnisch

et al. (1999) show that the growth rate of CF4 has slowed

in recent years from 1.07 ± 0.14 ppt yr-1 over the period

1978-1990 to 0.72 ± 0.11 ppt yr-1 between 1992 and 1998

(Figure 1-21). These growth rates correspond to global

emission rates of (16.3 ± 2.0) ¥ 106 kg yr-1 and (11.0 ±

1.7) ¥ 106 kg yr-1 and, allowing for estimated emissions

of CF4 from the semiconductor industry, imply a reduction of CF4 emissions from aluminum production of about

40%. This observation concurs with an aluminum

industry estimate of a 30% decrease in FC emissions

between 1990 and 1993 (Marks, 2000).

The current background atmospheric concentration

of C2F6 is about 3 ppt (Harnisch et al., 1996b; Oram, 1999;

Culbertson et al., 2000). Analysis of the Cape Grim air

archive shows that the mixing ratio of C2F6 in the Southern

Hemisphere has increased from 0.7 ppt in 1978 to 2.2 ppt

in late 1995 (Figure 1-21; Oram, 1999). The growth rate

over this period was approximately constant at 0.08 ppt

yr-1, which contrasts strongly with the behavior of CF4

whose growth rate has declined since the early 1990s.

Assuming emissions of C2F6 from aluminum production

have fallen along with those of CF4, an additional,

increasing source of C2F6 is required to maintain a constant

emission rate.

IPCC (Prather and Ehhalt et al., 2001) calculated

future CF4 and C2F6 abundances from 2000 to 2100 based

on respective emission rates in the six illustrative scenarios in the IPCC Special Report on Emissions Scenarios

(Nakic ́enovic ́ et al., 2000; Figure 1-21). Mixing ratios

projected for 2100 in the range of IPCC scenarios are 210-

400 ppt for CF4 (135-185 ppt in 2050) and 10-25 ppt for

C2F6 (7-10 ppt in 2050).

1.7.7.2 C-C4F8 AND C3F8

Since the previous Assessment (WMO, 1999), two

other FCs have been detected in the background atmosphere, perfluorocyclobutane (c-C4F8; FC-318) and perfluoropropane (C3F8; FC-218). The origin of c-C4F8 is

uncertain, although it is used in small amounts by the

semiconductor industry (cleaning and plasma etching),

and in very small amounts in some refrigeration blends

(Harnisch, 2000). C3F8 is used for plasma etching and a

small amount (~0.1 ¥ 106 kg yr–1) is emitted during aluminum production (Harnisch, 1997, 2000).

The abundance of c-C4F8 at Cape Grim increased

from 0.6 ppt in 1978 to about 1 ppt by late 1995 (Table 1-

12; Figure 1-21; Oram, 1999). The 1995 growth rate of

0.012 ppt yr–1 was about 80% lower than the maximum of

0.057 ppt yr–1 measured over the period 1983-1989. In

this way, its behavior is similar to that of CF4, whose

growth rate has slowed in response to reduced emissions

SOURCE GASES

1.59
from the aluminum industry. However, c-C4F8 has not

been reported in studies of the composition of emissions

from aluminum plants, so the cause of the observed

decrease in growth rate remains unclear. Measurements

of c-C4F8 have also been reported in stratospheric air.

Sturges et al. (2000a) reported a Northern Hemispheric,

upper tropospheric mixing ratio of about 1.0 ppt for 1994,

and Travnicek (1998) reported a 1997 mixing ratio of 0.7

ppt and calculated an annual global emissions rate of about

0.7 ¥ 106 kg yr–1 over the period 1978-1997.

In recent years the first atmospheric measurements

of C3F8 have also been reported. From a stratospheric

profile collected in 1997, Sturges et al. (2000a) estimated

an upper tropospheric mixing ratio of approximately 0.3

SOURCE GASES

1.60

5

1 0

1 5

2 0

2 5

ppt

1980 2000 2020 2040

year

B 2

A1

B 1

A 2

3

6

9

12

15

18

ppt

1980 2000 2020 2040

year

B2p

B1p

A2p

A1p

2

4

6

8

ppt

1980 2000 2020 2040

year

B2

B1

A1,A2

70

100

130

160

ppt

1980 2000 2020 2040

year

B 2

B1

A 2

A1

FC-318

FC-14 FC-116

sulfur hexafluoride

Figure 1-21. Long-term trends (annual mean data) and possible future levels of fluorocarbons and sulfur

hexafluoride. Data are from the Northern Hemisphere (pink line: MPAE, Harnisch et al., 1999, FC-14 only),

means from Cape Grim, Tasmania (blue lines: UEA, Oram, 1999; black line: SIO, Vollmer and Weiss, 2002,

SF6 only; orange line: UH, Maiss and Brenninkmeijer, 1998, SF6 only), and global surface means (red line:

CMDL, Hall et al., 2002, SF6 only). The scenarios (plotted from 2000) are from IPCC (2001; Appendix II).
ppt. The full profile between 11 and 32 km implies a long

atmospheric lifetime for this gas and recent tropospheric

increases. In addition, Culbertson et al. (2000) measured

a similar mean C3F8 mixing ratio of 0.26 ± 0.03 ppt in

samples collected at a rural, continental U.S. site in

March-April 2000.

1.7.7.3 SULFUR HEXAFLUORIDE (SF6)

Sulfur hexafluoride (SF6) is used mostly as an insulating fluid in step-down transformers and as a cover gas

in magnesium production. Atmospheric SF6 continues to

rise (Table 1-12, Figure 1-21), with current (2000) global

surface means estimated at 4.7 ppt (Hall et al., 2002) and

global growth rates of 4-5% yr–1. Vertical column abundances of SF6 monitored above the Jungfraujoch increased

by an average of 3.49 ¥ 1012 molec cm–2 yr–1 from 1987

to 2001, which corresponds to 12.9% yr–1 during 1987-

1988 and 5.1% yr–1 during 1999-2000 (Zander et al.,

2002). IPCC (Prather and Ehhalt et al., 2001) calculated

future SF6 abundances from 2000 to 2100 based on

respective emission rates in the six illustrative scenarios

in the IPCC Special Report on Emissions Scenarios

(Nakic ́enovic ́ et al., 2000; Figure 1-21). By 2100 mixing

ratios are estimated to be about 55-65 ppt (20-25 ppt in

2050) in all these scenarios.

1.7.7.4 TRIFLUOROMETHYLSULFURPENTAFLUORIDE

(SF5CF3)

Another long-lived greenhouse gas has recently

been found in the atmosphere (Sturges et al., 2000b).

Trifluoromethylsulfurpentafluoride (SF5CF3) has an

atmospheric lifetime of about 800 years (Table 1-3;

Takahashi et al., 2002) and, on a per-molecule basis, has

the largest radiative forcing of any gas found in the atmosphere to date. Measurements of SF5CF3 in air trapped in

polar snow (firn) show its abundance has risen from near

zero in the 1960s to about 0.12 ppt in 1999, and is currently growing at 0.008 ppt yr–1. The only known source

of SF5CF3 is as a byproduct of the manufacture of certain

fluorinated chemicals (Santoro, 2000).

1.8 HALOGENATED SOURCE GASES IN

THE FUTURE

Scenarios were developed to describe the range of

future mixing ratios of ODSs given our current understanding of sources, sinks, banks or reservoirs of unemitted material and emission rates from these, and potential future production. The baseline scenario Ab was

designed as a best estimate of future trace gas mixing

ratios based upon current trends in production and consumption, and limits to these as defined by the fully

amended and adjusted Montreal Protocol. Scenario Am

was an attempt to define atmospheric mixing ratios in a

world where future production and consumption rates

were at the maximum allowed by the fully amended and

adjusted Montreal Protocol. It must be kept in mind that

for some gases, notably the HCFCs, this concept is ill

defined because upper limits to global production apply

only after 2016. The Pc scenario was constructed to show

how continued constant production of ODSs at 1999 rates

into the future, in clear violation of the Montreal Protocol,

would affect future EESC. Hypothetical scenarios were

also developed in which emissions of all ODSs (E0) or a

selected class of ODSs ceased in 2003. These hypothetical cases were constructed to show how future emissions

of different compound classes affect future EESC and to

define a lower limit to future EESC. Additional, potentially achievable cases were also constructed to show the

effects of zero production from 2003 onward for groups

of ODSs and all ODSs together (P0). More details

regarding the assumptions made in developing these scenarios and cases are included in Tables 1-13 and 1-14.

1.8.1 Development of the Baseline

Scenario Ab

To create the baseline scenario Ab, a set of emissions was derived for each halocarbon for 1950-2000 so

that mixing ratios calculated in a simple box model would

match as closely as possible the observed mixing ratios

(Tables 1-15 and 1-16; Figure 1-22). Where possible,

emissions were based on historical industry production

data and emission functions as described in Section 1.3

(e.g., McCulloch and Midgley, 2001; McCulloch et al.,

2001; AFEAS, 2001; UNEP, 2002). Some discrepancies

exist, however, between measured and calculated mixing

ratios (see Section 1.6), and in order to provide a history

without discontinuities, scale factors were sometimes

applied to these emissions to better match the observed

atmospheric mole fractions (Tables 1-15 and 1-16).

Future emissions were derived for 2000-2100

based upon (1) current trends in halocarbon production

and emission, (2) the size of unemitted reserves (adjusted

to reflect scale factors applied to past emissions), (3) relationships between production and emission as a function

of application, and between reserves and emission, and

(4) considering future production restrictions outlined in

the latest Amendments (Beijing) of the Montreal Protocol.

The Beijing Amendments (1999) to the Montreal Protocol

included limits on the production of HCFCs in both developed (freeze in 2004) and developing (freeze in 2016)

countries, stricter limits on the production of ODSs by

SOURCE GASES

1.61
developed countries for use in developing countries, as

well as a global phaseout of a new species, bromochloromethane (CH2BrCl), in 2002.

For methyl chloride and methyl bromide, scenarios

were developed for past years based upon available firn

air and real-time measurements (Khalil et al., 1993; Khalil

and Rasmussen, 1999; Butler et al., 1999; Sturges et al.,

2001a). Mixing ratios were set for the late 1990s at 9.5

ppt CH3Br and 550 ppt for CH3Cl based upon discussions

in Kurylo and Rodríguez et al. (1999). Future mixing ratios

in scenarios were calculated assuming that restrictions on

production resulted in emission reductions in the same year

(Table 1-13). Future emissions of CH3Cl were kept constant to sustain a mixing ratio of 550 ppt for all scenarios.

1.8.2 Projecting Future Halocarbon

Mixing Ratios

Future mixing ratios of a trace gas depend upon

future emission rates and the rate at which the gas is

removed from the atmosphere. Atmospheric removal

rates are determined by the chemical and physical characteristics of a trace gas, and cannot be purposefully altered

by humans. In the absence of emissions, future mixing

ratios can be calculated on the basis of trace gas lifetimes.

Because lifetimes of ODSs range from less than 1 year to

more than 100 years, however, the expected rate of decline

in the absence of emissions varies substantially for different halocarbons.

Emissions of most ODSs, however, are not yet

insignificant. Substantial amounts of some ODSs exist in

current applications and will be released to the atmosphere in future years. Furthermore, production of some

halocarbons continues under the allowances prescribed

by the Montreal Protocol. Predicting future mixing ratios

of ODSs requires a good understanding of emission rates

of ODSs from present-day applications, the size of produced but unemitted reserves (banks), the magnitude of

future production, and halocarbon atmospheric lifetimes.

Managing or controlling future mixing ratios is possible

only through controlling current rates of emission and

future production.

In what follows, some salient points regarding

future halocarbon mixing ratios and the development of

scenarios are discussed.

1.8.2.1 CFCS

Reserves of unemitted but produced CFCs vary

substantially for different individual CFCs because they

were used in different applications. CFC-11 was used

extensively in foams from which only a small portion of

the CFC becomes released each year. As a result a very

large reservoir of CFC-11 exists in foams today, and future

emissions will depend mostly upon the rate at which CFC11 escapes from foams together with the amount of future

production. Emission models suggest that the reservoir

of unemitted CFC-12 and -113 is substantially smaller,

owing to use in applications where release is more rapid

(AFEAS, 2001). For these two gases, future emissions

will be more closely tied to the magnitude of future production. Differences in the calculated future scenarios

reflect the size of reserves for each gas relative to the

potential magnitude of future production (Figure 1-1).

1.8.2.2 METHYL CHLOROFORM

Reserves of produced but unemitted methyl chloroform (CH3CCl3) are thought to be small owing to past

use predominantly as a solvent and cleaning agent.

Accordingly, future emissions should be closely tied to

future production. Known production in recent years has

been about 20 Gg yr–1 (McCulloch and Midgley, 2001),

or quite small compared with earlier years, and should

decrease as the allowances for production and consumption in developing countries diminish through 2015. The

exponential decline in mixing ratios in recent years will

continue as long as methyl chloroform emissions remain

small relative to its atmospheric burden. Future production within limits allowed by the Protocol could slow this

decline (Figure 1-4). Furthermore, if some nonindustrial

emissions exist and persist in the future, perhaps as a result

of biomass burning (Rudolph et al., 2000), the exponential rate of decay would slow until some non-zero steadystate mixing ratio was reached. Current estimates of nonindustrial emission magnitudes suggest that this steadystate mixing ratio would be <2 ppt (see Section 1.3). 

1.8.2.3 CARBON TETRACHLORIDE

Significant mismatches between observed and calculated mixing ratios of carbon tetrachloride (CCl4) suggest potential difficulty in reliably estimating future

amounts of this important ODS (see Section 1.6).

Improvements in this area will come only after certain

questions are addressed. These questions include the following: (1) How much carbon tetrachloride is emitted

from processes not related to the production of CFCs, and

how have these releases changed over time? (2) Are fugitive releases of CCl4 on the order of 11% from production

of CFCs in the years 1978-1993 reasonable? (3) Did those

fugitive releases more than double to become about 25%

in 1995-1999 as the bulk of CFC production shifted to

countries operating under Article 5(1) of the Montreal

Protocol (Figure 1-15; Article 5(1) refers to developing

countries as defined by the Montreal Protocol)?

SOURCE GASES

1.62
SOURCE GASES

1.63

Table 1-13. Description of scenarios for future ODS emissions.

Ab:

Baseline 

(best-guess

scenario

following the

Beijing

Amendments,

1999) =

Reference

Am:

Maximum

production

Production of CFCs, halons, and HCFCs is projected forward at the maximum allowed under the

Protocol.

CFCs: The maximum global production allotment for CFCs is, according to UNEP (2002), 160 ODP-ktons

yr–1 through 2004, decreasing to zero in 2010. This is 17.5% above the 1999 production in the Ab scenario.

For the maximum scenario, the Ab production of CFC-11, -12, and -113 was therefore multiplied by 1.175

for the years 2000-2004 (a scale factor of 1.0 was applied to CFC production during 2005-2009). The emissions of carbon tetrachloride were consequently also multiplied by the same factor, since they are linked in

the calculation to the production of CFC-11 and -12. The UNEP production limit for CCl4 was not applied.

Halons: For Halon-1211 a global bank of 114 Gg in 2000 was used, instead of 98 Gg as in Ab. For the combined halons, a total global production was applied of 45 ODP-ktons through 2004, and 22.5 ODP-ktons split

among the years 2005-2009. This production was divided over Halon-1211 and -1301 according to their production split in 1999 in the Ab scenario.

HCFCs: Future production and consumption of HCFCs was estimated by keeping production in non-Article

5(1) countries constant at the 1999 level of 27 ODP-ktons yr–1 and by increasing production in Article 5(1)

countries by 3% yr–1 starting in 1999 at 9.4 ODP-ktons yr–1. The Protocol limits for production and consumption were applied for later years. The production of HCFCs in future years was split among HCFC-22,

-141b, and -142b, as it was in 1999 (AFEAS, 2001). Production of methyl chloroform and methyl bromide

are all at their practical maxima and so are the same as in the Ab scenario.

Developed countries (non-Article 5(1)):

Future production of CFCs was continued at the current level of capacity for "essential uses" and export to

Article 5(1) countries. The bank of unemitted reserves in 2000 was estimated to be 656 Gg for CFC-11 and

60 Gg for CFC-12. Future production of CCl4 was governed by the quantities required for CFC manufacture;

global emissions of CCl4 were estimated as a fraction (40%) of global CFC production (see Figure 1-15).

Future production of methyl chloroform was zero. Emission of halons in future years was estimated as a percentage of unemitted global reserves: 98 Gg for Halon-1211 and 62 Gg for Halon-1301 in 2000. Production

of HCFCs was continued with the same trends as now; after 2003 production is assumed to decline according

to consumption limits required by the Protocol. Controls on global production of CH3Br were presumed to

result in a global emission decrease of 30 Gg yr–1 (out of 212 Gg yr–1) in the future as governed by the phaseout schedule of the Protocol; 75% of this emission was phased out according to the schedule for production

decreases in non-Article 5(1) countries, and 25% was phased out according to this schedule in Article 5(1)

countries. Emissions from quarantine pre-shipment applications were held constant into the future at 12 Gg

yr–1; emissions from biomass burning and gasoline combustion were also held constant in all scenarios.

Developing countries (Article 5(1)):

Production of CFCs and methyl chloroform was allowed to continue at the current (frozen) level, with phase

down dates mandated in the Montreal Protocol, and there was no production allotted in Countries with

Economies in Transition (the de facto situation). Emissions of carbon tetrachloride were estimated at almost

40% of CFC-11 and -12 produced. Emissions of halons were estimated as a percentage of unemitted global reserves. The size of this unemitted stock of Halon-1211 in 2000 was enhanced slightly to account for

potential additional production in Article 5(1) countries. For Halon-1301 production in A5(1) countries was

taken from Fraser et al. (1999). Production and consumption of HCFCs was kept constant at current levels

until phase down is required (after 2015). For a discussion of future production and emission of CCl4 and

CH3Br, see the discussion for non-Article 5(1) countries.

Scenario Description

Continued production of all ODSs at a constant level reported by UNEP (2002) for the year 1999 was

assumed. This additional production would be in excess of the limits set by the amended Montreal Protocol.

Global production magnitudes reported by UNEP (ODP-Gg yr–1) for 1999 and used in this scenario were:

147.2 for CFCs, 25.2 for halons, 1.78 for methyl chloroform, and 36.2 for HCFCs. Carbon tetrachloride production was kept constant in the future at 67.7 Gg yr–1 based upon production of CFC-11 and -12 in 1999.

Methyl bromide emission was held constant at the 1999 level. For the CFCs (CFC-11, -12, and -113), halons

(1211 and 1301), and HCFCs (HCFC-22, -141b, and -142b), the 1999 production of the individual species

was calculated from the aggregated ODP-weighted UNEP total using the ratio in the Ab scenario for 1999.

Pc:

Continued

production
SOURCE GASES

1.64

Table 1-14. Description of cases for future ODS production and emissions.

Case Description

E0: Zero emission

(hypothetical scenario)

Anthropogenic emissions of all ODSs set to zero from 2003 onward, i.e., banks destroyed and

production stopped. In practice, this is an unachievable scenario. This case is used to show the

natural removal rates of ODS from the atmosphere and describes a lower limit for ODS mixing

ratios in the future atmosphere.

Emission contribution

of CFCs to Ab

No emission of CFCs (including from existing equipment) from 2003 onward. Other species as

in scenario Ab.

Emission contribution

of CCl4 to Ab

No emission of carbon tetrachloride (including from existing equipment) from 2003 onward.

Other species as in scenario Ab. Since the emissions of CCl4 are fugitive and in the calculations

related to the production of CFC-11 and -12, this is the same as a zero production case.

Emission contribution

of CH3CCl3 to Ab

No emission of methyl chloroform (including from existing equipment) from 2003 onward. Other

species as in scenario Ab.

Emission contribution

of halons to Ab

No emission of halons (including from existing equipment) from 2003 onward. Other species as

in scenario Ab.

Emission contribution

of HCFCs to Ab

No emission of HCFCs (including from existing equipment) from 2003 onward. Other species as

in scenario Ab.

Emission contribution

of CH3Br to Ab

No anthropogenic emission of methyl bromide (including from existing equipment) from 2003

onward. Total reduction of 42 Gg (30 for controllable uses and 12 for Quarantine and PreShipment (QPS) uses). Other species as in scenario Ab. The emissions of CH3Br are from dispersive uses, so there is no substantial bank. As a result, this case is identical to a zero methyl bromide production case. Nonindustrial emissions of 160-170 Gg yr–1 were presumed and were kept

constant in the future.

P0: Zero production Production of all ODSs set to zero from 2003 onward. Emissions continue based upon size of

unemitted reserves (based on the cases described below).

Production contribution of CFCs to AbNo production of CFCs from 2003 onward. Emission of CFC-12 equals its production after 2002

(bank is depleting rapidly), so for CFC-12 this equals zero emissions. Production of CFC-114

and -115 already zero in Ab. Non-CFCs as in scenario Ab.

Production contribution of CH3CCl3 to AbNo production of methyl chloroform from 2003 onward. Other species as in scenario Ab.

Production contribution of halons to AbNo production of halons from 2003 onward. Global production of Halon-1211 and -1301 in

1999 is, according to Fraser et al. (1999), 29.0 ODP-ktons and according to UNEP (2002), 25.2

ODP-ktons in 1999. This case was derived by estimating emissions from the halon production

predicted in Fraser et al. (1999) for 2003-2009 (zero after 2009) and subtracting these emissions

from those calculated in the Ab scenario. Halon-1202 emissions were presumed to be zero, since

they arise as a byproduct of Halon-1211 production. Halon-2402 production was limited to

Russia. The amount produced decreased rapidly from 2.45 ktons in 1990 to 0.152 ktons in 1999

(UNEP, 1999). The 1996 number is only a small fraction of the emission. Halon-2402 production

is therefore assumed zero in this case (from 2003 onward) (same as in Ab, in agreement with

Madronich and Velders et al. (1999) and Fraser et al. (1999)).

Production contribution of HCFCs to AbNo production of HCFCs from 2003 onward. Other gases as in scenario Ab.

Production contribution of CH3Br to AbNo production of CH3Br from 2003 onward. Because emissions are presumed to be prompt after

production, this case is identical to the case of zero CH3Br emissions.
SOURCE GASES

1.65

Table 1-15. Source of production, emission, and observation data used in creating scenarios.

Details Related to Creating Mean Details Related to Calculating Expected Mixing Ratios

Set of Global Observations a and Past Mixing Ratios in Scenarios from

Mean Period Source of Mean a Simple Box Model Calculation b

AGAGE/of Observations Lifetime Production Emission Emission

CMDL Ratio Overlap after Scaling c Used In Data Function Scale

(s.d.) Box Model Source Revised? Factor d

(years)

CFC-12 1.008 1991-1999 AGAGE 100 AFEAS, No 1.045 ( 1980)

(0.002) UNEP 0.8 (1993-94)

CFC-11 0.980 1990-1999 AGAGE 45 AFEAS, Yes, foam e 1.068

(0.002) UNEP

CFC-113 0.993 1994-1999 AGAGE 85 AFEAS No 0.98

(0.005) only

CCl4 0.973 1993-1999 AGAGE 26 ff f

(0.006)

CH3CCl3 0.975 1993-1999 AGAGE 4.8 g No Variable

(0.008) (0.9 to 1.25)

HCFC-22 0.993 1998-2000 CMDL 11.8 AFEAS, No 1.0

(0.007) UNEP

HCFC-141b 1.011 1998-2000 CMDL 9.2 AFEAS, No 1.0

(0.006) UNEP

HCFC-142b 1.049 1998-2000 CMDL 18.5 AFEAS, Yes, foam h 1.0

(0.008) UNEP

Halon-1211 1.068 1998-2000 CMDL 16 i No 1.18 ( 1992)

(0.006)

Halon-1301 1.139 1998-2000 CMDL 65 j No Variable

(0.004)

CH3Br NA NA k 0.7 NA NA NA

CH3Cl NA NA l 1.3 NA NA NA

NA, Not available.

a Mean halocarbon measurements were derived from the longest available global data record after adjusting this data record by the factor [1 ± (1–x)/2],

where x is the mean AGAGE/CMDL ratio indicated. Scenarios were created to match these measurements for past years. AGAGE data are reported

on the SIO-98 scale (Prinn et al., 2000); CMDL data are reported on scales as described in Hall et al. (2002).

b Calculated mixing ratios also appear in Scenarios and provide the basis for calculating future emissions and mixing ratios. Citations referred to are

AFEAS (2001) and UNEP (2002) unless otherwise specified. c Data record from the group indicated was scaled by half the AGAGE vs. CMDL ratio.

d To create a set of emissions consistent with global observations (Section 1.8 only), emissions derived from industrial production were scaled by this

factor up to the point at which measurements become available. e McCulloch et al. (2001).

f CCl4 emission estimated in the future from CFC-11 and -12 production (see text). g McCulloch and Midgley (2001).

h Emission functions for foams adjusted as described in Ashford (2000) and McCulloch et al. (2001).

i Updated UNEP (1999).

j Fraser et al. (1999).

k Southern Hemispheric firn data were scaled by a constant of 1.16 to approximate global mixing ratios (see Section 1.5).

l Southern Hemispheric firn data were scaled by a constant of 1.06 to approximate global mixing ratios (see Section 1.5).
SOURCE GASES

1.66

Table 1-16. Calculated global surface mixing ratios (ppt) in the Ab baseline scenario. Dates are for the beginning

of the year indicated.

Year CFC-12 CFC-114 CCl4 HCFC-22 HCFC-142b Halon-1202 Halon-2402 CH3Cl

CFC-11 CFC-113 CFC-115 CH3CCl3 HCFC-141b Halon-1211 Halon-1301 CH3Br

1990 256.3 474.1 68.0 15.4 5.1 105.7 124.7 82.0 0.0 1.2 2.27 0.03 1.66 0.31 9.0 550.5

1991 262.6 488.0 74.0 15.7 5.6 105.8 128.8 88.1 0.0 1.7 2.54 0.03 1.85 0.34 9.0 550.6

1992 267.4 500.2 78.3 15.9 6.0 105.5 129.1 94.4 0.0 2.4 2.79 0.03 1.98 0.36 9.1 550.6

1993 271.2 511.5 81.4 16.0 6.5 105.1 127.8 100.8 0.2 3.3 2.94 0.03 2.07 0.38 9.2 550.6

1994 271.2 518.7 82.9 16.1 7.0 104.5 118.8 106.8 0.9 4.2 3.10 0.04 2.17 0.40 9.3 550.7

1995 270.6 524.8 83.2 16.2 7.5 103.7 107.8 112.8 2.1 5.2 3.25 0.04 2.33 0.42 9.4 550.7

1996 269.5 530.0 83.2 16.2 7.9 102.7 97.0 118.6 3.8 6.3 3.40 0.04 2.46 0.43 9.5 550.7

1997 268.2 534.2 82.7 16.3 8.3 101.6 82.9 125.0 5.5 7.4 3.55 0.05 2.55 0.43 9.5 550.7

1998 266.6 537.4 82.6 16.3 8.6 100.5 69.3 130.5 7.1 8.5 3.69 0.05 2.64 0.43 9.5 550.7

1999 264.7 539.5 82.3 16.3 8.8 99.2 58.6 136.1 9.0 9.4 3.83 0.05 2.73 0.42 9.5 550.7

2000 262.6 540.5 81.7 16.4 9.0 97.9 49.9 141.6 10.8 10.5 3.97 0.05 2.82 0.41 9.3 550.7

2001 260.5 541.2 81.1 16.4 9.1 96.7 42.6 149.3 13.0 12.2 4.07 0.04 2.90 0.41 9.3 550.7

2002 258.2 541.9 80.4 16.4 9.1 95.5 36.6 156.4 15.4 13.9 4.13 0.04 2.97 0.40 9.1 550.7

2003 255.8 542.7 79.6 16.4 9.1 94.4 31.6 162.9 18.1 15.5 4.16 0.03 3.04 0.39 9.0 550.7

2004 253.3 541.9 78.8 16.4 9.2 93.3 27.5 168.9 20.9 17.1 4.16 0.03 3.11 0.38 8.9 550.7

2005 250.7 541.1 78.0 16.4 9.2 92.3 24.0 174.4 23.8 18.7 4.14 0.02 3.17 0.37 8.8 550.7

2006 248.0 538.5 77.2 16.4 9.2 90.3 21.2 178.2 26.5 20.1 4.09 0.02 3.23 0.36 8.5 550.7

2007 245.2 536.0 76.3 16.4 9.2 88.5 18.8 180.6 28.9 21.3 4.03 0.02 3.28 0.35 8.5 550.7

2008 242.4 531.4 75.5 16.3 9.2 85.6 16.7 182.0 31.1 22.3 3.95 0.01 3.32 0.34 8.4 550.7

2009 239.3 526.9 74.7 16.3 9.2 82.8 15.0 182.7 33.0 23.2 3.86 0.01 3.37 0.32 8.4 550.7

2010 236.1 522.5 73.9 16.3 9.2 80.1 13.6 182.8 34.8 24.0 3.76 0.01 3.40 0.31 8.4 550.7

2011 232.7 517.3 73.0 16.2 9.2 77.1 12.3 182.5 36.5 24.7 3.66 0.01 3.44 0.30 8.4 550.7

2012 229.2 512.2 72.2 16.2 9.2 74.2 11.2 180.8 37.7 25.2 3.54 0.00 3.46 0.29 8.4 550.7

2013 225.6 507.1 71.3 16.2 9.2 71.4 10.3 178.2 38.5 25.5 3.43 0.00 3.49 0.28 8.4 550.7

2014 221.9 502.0 70.5 16.1 9.2 68.7 9.5 174.8 39.1 25.6 3.31 0.00 3.51 0.27 8.4 550.7

2015 218.2 497.0 69.7 16.1 9.2 66.1 8.7 171.0 39.4 25.7 3.19 0.00 3.52 0.26 8.4 550.7

2016 214.3 492.1 68.9 16.0 9.1 63.6 8.0 167.0 39.6 25.7 3.07 0.00 3.54 0.25 8.2 550.7

2017 210.5 487.2 68.1 16.0 9.1 61.2 7.4 162.0 39.3 25.5 2.95 0.00 3.55 0.24 8.2 550.7

2018 206.6 482.3 67.3 15.9 9.1 58.9 6.9 156.2 38.8 25.1 2.83 0.00 3.55 0.23 8.2 550.7

2019 202.7 477.5 66.5 15.9 9.1 56.7 6.3 150.1 38.0 24.7 2.71 0.00 3.56 0.22 8.2 550.7

2020 198.9 472.8 65.7 15.8 9.1 54.6 5.9 143.9 37.1 24.1 2.59 0.00 3.56 0.21 8.2 550.7

2021 195.0 468.1 64.9 15.8 9.1 52.5 5.4 137.6 36.1 23.6 2.48 0.00 3.55 0.20 8.2 550.7

2022 191.2 463.4 64.2 15.7 9.1 50.5 5.0 131.2 34.9 22.9 2.37 0.00 3.55 0.19 8.2 550.7

2023 187.4 458.8 63.4 15.7 9.1 48.6 4.7 124.8 33.7 22.3 2.26 0.00 3.54 0.18 8.2 550.7

2024 183.6 454.2 62.7 15.6 9.1 46.8 4.3 118.6 32.4 21.6 2.16 0.00 3.53 0.18 8.2 550.7

2025 179.9 449.7 62.0 15.6 9.1 45.0 4.0 112.6 31.0 20.9 2.05 0.00 3.52 0.17 8.2 550.7

2030 162.0 427.8 58.4 15.3 9.1 37.1 2.8 86.8 24.7 17.6 1.59 0.00 3.44 0.13 8.2 550.7

2035 145.5 406.9 55.1 15.1 9.0 30.6 1.9 66.9 19.1 14.6 1.22 0.00 3.33 0.10 8.2 550.7

2040 130.5 387.1 51.9 14.8 9.0 25.3 1.3 51.6 14.6 12.0 0.93 0.00 3.19 0.08 8.2 550.7

2045 116.9 368.2 49.0 14.6 9.0 20.9 0.9 39.3 11.0 9.8 0.70 0.00 3.04 0.06 8.2 550.7

2050 104.7 350.2 46.2 14.4 9.0 17.2 0.6 27.6 7.7 7.6 0.52 0.00 2.88 0.05 8.2 550.7

2055 93.8 333.2 43.5 14.1 8.9 14.2 0.5 18.7 5.2 5.8 0.39 0.00 2.72 0.04 8.2 550.7

2060 83.9 316.9 41.0 13.9 8.9 11.7 0.3 12.5 3.4 4.4 0.29 0.00 2.56 0.03 8.2 550.7

2065 75.1 301.5 38.7 13.7 8.9 9.7 0.2 8.3 2.1 3.4 0.21 0.00 2.41 0.02 8.2 550.7

2070 67.2 286.8 36.5 13.4 8.9 8.0 0.2 5.5 1.3 2.6 0.16 0.00 2.25 0.02 8.2 550.7

2075 60.2 272.8 34.4 13.2 8.8 6.6 0.1 3.6 0.8 1.9 0.12 0.00 2.11 0.01 8.2 550.7

2080 53.8 259.5 32.4 13.0 8.8 5.4 0.1 2.4 0.5 1.5 0.09 0.00 1.97 0.01 8.2 550.7

2085 48.2 246.8 30.6 12.8 8.8 4.5 0.1 1.6 0.3 1.1 0.06 0.00 1.83 0.01 8.2 550.7

2090 43.1 234.8 28.8 12.6 8.8 3.7 0.0 1.0 0.2 0.8 0.05 0.00 1.71 0.01 8.2 550.7

2095 38.6 223.3 27.2 12.4 8.7 3.1 0.0 0.7 0.1 0.6 0.03 0.00 1.59 0.01 8.2 550.7

2100 34.5 212.4 25.6 12.2 8.7 2.5 0.0 0.5 0.1 0.5 0.03 0.00 1.48 0.00 8.2 550.7
In the Ab and Am scenarios, future emissions of

CCl4 were estimated from CFC production and a molar

ratio [total CCl4 emission]/[global CFC production] of

0.38 as estimated for recent years (see Figure 1-15). These

projections (Figure 1-4) will be reasonable for carbon

tetrachloride only if all significant emissions of this gas

become diminished in the future along the time frame projected for reductions in CFC production.

1.8.2.4 HALONS

A large fraction of present-day halon emissions

arise from use of a banked halon, defined here as material

that is contained in inventories and installed equipment.

As a result, future mixing ratios are tied to the rate of emissions from these halon banks in addition to the magnitude

of future production (Figure 1-3). Although progress has

been made in reducing production for future uses (UNEP,

1997), halon banks are large relative to current emission

rates (Butler et al., 1999; Fraser et al., 1999; UNEP, 1999,

and updates). In the Ab scenario, mixing ratios of Halon1301 continue to increase until about 2020 primarily

because of emission from reserves. Scenario results for

Halon-1211 also suggest that emissions from reserves are

important, but the magnitude and date of peak mixing

ratios are strongly influenced by future production rates

as well (Figure 1-3).

Additional uncertainty in future halon mixing ratios

arises because the size of these reserves is not well

defined; this added uncertainty is not explicitly included

in the calculated scenarios but is considered elsewhere

(see Section 1.8.4). For Halon-1211, estimates of reserves

derived from production data and functions relating

releases to production vary from 65 to 136 Gg in 2000

(Fraser et al., 1999; UNEP, 1999, updated). Reserves calculated with emissions derived from atmospheric measurements (Butler et al., 1998, 1999) and an atmospheric

lifetime of 16 years are about 100 Gg in 2000; those

SOURCE GASES

1.67

400

600

500

300

200

100

CFC-12

1950 2000 2050 2100 ppt 4.0

3.0

2.0

1.0

0.0

Halon-1301

Halon-2402

1950 2000 2050 2100 ppt ppt ppt

400

600

500

300

200

100

10

1950 2000 2050 2100

200

160

120

80

40

1950 2000 2050 2100

CH CCl

HCFC-22

HCFC-141b

HCFC-142b ppt

CFC-11

CFC-113

Halon-1211

CCl

CH Cl

CH Br

Figure 1-22. Global atmospheric surface mixing ratios (ppt) in the baseline scenario Ab (lines) and the mean

AGAGE-CMDL observations (circles; see Table 1-15 for origin; Prinn et al., 2000; Montzka et al., 1999). 
derived from independent measurements at Cape Grim

by UEA (Fraser et al., 1999) are substantially smaller. For

reference, current increases observed for Halon-1211 of

0.1 ppt yr–1 would require emissions of about 10 Gg yr–1

to sustain (for a 16-yr lifetime). For Halon-1301, estimates of reserves range from 45 Gg (UNEP, 1999,

updated) to 65 Gg (Fraser et al., 1999) and, for reference,

measurements suggest annual emissions of about 3 Gg

yr–1 (for a 65-yr lifetime).

1.8.2.5 HCFCS

Uncertainty in future mixing ratios of HCFCs stems

primarily from projecting how these gases will fulfill the

needs of developing nations. Consumption is currently

frozen in non-Article 5(1) countries, and soon reductions

in consumption will be mandated there as well (see Figure

1-18). Production in non-Article 5(1) countries, however,

is capped only after 2004; this production could provide a

substantial allotment of HCFC for use in Article 5(1)

countries, depending upon demand. In addition, production in Article 5(1) countries is not restricted until after

2015, although the extent to which these countries would

invest substantial capital in the manufacture and use of

such time-limited materials remains a major uncertainty.

Production (as ODP-tons) in developing countries averaged 10% of production in non-Article 5(1) countries in

the late 1990s (UNEP, 2002).

In the Ab scenario, future global HCFC production

was presumed to decrease according to the consumption

limits outlined for non-Article 5(1) countries and a small,

constant allotment was included for production in countries operating under Article 5(1) of the Protocol. This

scenario would underestimate future production, emissions, and consequent mixing ratios of HCFCs if demand

for these gases in developing countries increased substantially in future years. In the Am scenario, production and

consumption of HCFCs was increased as described in

Table 1-13. This enhanced demand leads to higher HCFC

mixing ratios in the future (Figure 1-5); this enhanced

HCFC production accounts for about one-third of the

14.5% higher integrated EESC calculated for scenario Am

compared with scenario Ab (Table 1-17).

1.8.2.6 METHYL BROMIDE AND METHYL CHLORIDE

Atmospheric releases resulting from industrial

production represent only a fraction of total releases of

these gases to today’s atmosphere (see Section 1.5).

Uncertainty in the magnitude of this fraction limits our

ability to accurately describe how methyl bromide mixing

ratios will respond to production limits in the revised and

amended Montreal Protocol (Section 1.5.1.4). Scenarios

for methyl bromide were developed by prescribing future

emission decreases of about 30 Gg, or 15% of the total

annual flux, according to the schedule of production

restrictions in the current Protocol (Table 1-13; see also

Section 1.5). Anthropogenic emissions from quarantine

and preshipment applications (12 Gg yr–1) and from nonfumigation activity (biomass burning and automobile

emissions) were kept constant in the future in all scenarios.

No limits apply to methyl chloride production or consumption; total emissions of this gas were held constant

in the future in all scenarios.

Both methyl bromide and methyl chloride are

unique among ozone-depleting halocarbons because a

substantial fraction of their sources and sinks arises from

biological processes. These processes may be affected

significantly by changes in future climate. Because our

understanding of these process is limited, we cannot predict if such changes would act to enhance or diminish

mixing ratios of these gases in the future atmosphere (see

also Section 1.8.5).

1.8.2.7 EFFECTIVE EQUIVALENT CHLORINE

As the influence of methyl chloroform on effective

equivalent chlorine (EECl) has lessened, the decline

observed for EECl has slowed (see Section 1.2; Figure 1-

9). This is because mixing ratios of CFC-12, halons, and

the HCFCs continued to increase in 2000. Future trace

gas trends in scenario Ab (the best-guess scenario) result

in continued declines in EECl in the near future; the mean

rate of change for EECl during 2000-2010 in this scenario

is –0.9% yr–1 (Figure 1-9). In scenario Ab, EECl continues to decline because projected mixing ratios for CFC12 and total bromine peak and subsequently begin

decreasing during 2000-2005, and because enhanced

declines in other CFCs are projected for that period as

well (Figure 1-22). In scenario Am, EECl decreases at

–0.7% yr–1, slightly slower than in the Ab scenario.

1.8.3 Scenario Results: Calculated Mixing

Ratios and EESC

The concept of effective equivalent stratospheric

chlorine (EESC) has been used in the past to gauge the net

effect of tropospheric halocarbon trends on the loading of

ozone-depleting halogen in the stratosphere (Prather and

Watson, 1990; Solomon and Wuebbles et al., 1995;

Madronich and Velders et al., 1999; Daniel et al., 1995).

As such, EESC provides a rough estimate of the time scale

for ozone recovery in an otherwise unchanging atmosphere. The concept of EESC has limitations, however, particularly because of stratospheric mixing processes and

other effects (Prather, 1997; see also discussion in Section

SOURCE GASES

1.68
1.2.5). Furthermore, the chemical and dynamical state of

the atmosphere has not been constant, and changes in

aerosols, trace gas mixing ratios, and stratospheric temperatures are also expected to influence ozone chemistry

and the efficiency of halogens to deplete stratospheric

ozone in the future. Whereas different scenarios are

gauged in this chapter in terms of integrated EESC and the

date when EESC returns to the value estimated for 1980,

discussions relating to future ozone concentrations in light

of the broad range of additional atmospheric changes that

are possible can be found in Chapters 3 and 4.

EESC was calculated as described in past

Assessments for the updated scenarios. These calculations were performed with a slightly reduced efficiency

for bromine relative to chlorine, on a per-atom basis, for

the destruction of ozone compared to WMO (1999) (a =

45; see Section 1.4.4). Furthermore, an absolute fractional

release of 0.8 for CFC-11 was included in this calculation

to provide consistency with past reports, although the data

of Schauffler et al. (2002) suggest that a more appropriate

estimate may be 0.55. This absolute degradation amount

does not affect calculations of ODP or relative changes in

EESC presented in this chapter, but it is important when

considering the influence short-lived gases have on the

burden of ozone-depleting halogen, particularly in the

lower stratosphere (see Chapter 2).

Despite the limitations of EESC, different scenarios

and cases are compared with respect to the year (x) EESC

returns to the value estimated for 1980, and by the percentage differences in time-integrated EESC from 2002

or 1980 through the year (x) (Table 1-17; Solomon and

Wuebbles et al., 1995; Madronich and Velders et al.,

1999). Because of the revised value for alpha, and because

time-varying past histories for CH3Br and CH3Cl are now

SOURCE GASES

1.69

Table 1-17. Comparison of scenarios and cases: the year when EESC drops below the 1980 level, and

integrated EESC differences (relative to baseline scenario Ab). a

Scenarios and Cases Year (x) when Percent difference in Percent difference in

EESC is expected

to drop below 1980

value a

above the 1980 above the 1980

level relative to level relative to

scenario Ab a scenario Ab a

Scenarios

Ab: Beijing Amendments (1999) 2043.9 0.0 0.0

Am: Maximum allowed production 2048.0 8.3 14.5

Pc: Continued production Not reached >115 b >202 b

Cases c

P0: Zero production of all anthropogenic ODSs 2039.9 –8.1 –14.1

Contribution CFC production to Ab 2042.8 –2.2 –3.8

Contribution CH3CCl3 production to Ab 2043.9 –0.2 –0.3

Contribution Halon production to Ab 2043.6 –0.6 –1.1

Contribution HCFC production to Ab 2042.6 –3.1 –5.4

Contribution anthropogenic CH3Br production to Ab 2042.9 –2.3 –4.0

E0: Zero emission of all anthropogenic ODSs 2033.8 –19.3 –33.8

Contribution CFC emissions to Ab 2041.1 –5.4 –9.4

Contribution CCl4 emissions to Ab 2043.5 –1.5 –2.7

Contribution CH3CCl3 emissions to Ab 2043.7 –0.9 –1.6

Contribution Halon emissions to Ab 2040.7 –6.1 –10.8

Contribution HCFC emissions to Ab 2042.3 –4.9 –8.6

Contribution anthropogenic CH3Br emissions to Ab 2042.9 –2.3 –4.0

a The dates and percent differences in integrated EESC were calculated with the fractional release values in column 2 of Table 1-4.

b The integrated percent differences indicated for Scenario Pc represent lower limits because the integration was run only through 2100 and EESC in

this scenario does not reach the 1980 level by 2100 (see Figure 1-23). c Contributions of ozone-depleting substances to future EESC were calculated in the hypothetical “Cases” by setting production or emission to zero in

2003 and subsequent years for the indicated chemical or group of chemicals.

EESC 

1980

dt

x

Ú EESC 

2002

dt

x

Ú
considered (Section 1.8.1), the EESC value in 1980 is calculated here to be 1.784 ppb, somewhat lower than in

Madronich and Velders et al. (1999).

Future EESC and time-integrated EESC were calculated from halocarbon mixing ratios in selected scenarios (Table 1-17). For the baseline scenario (Ab), the

calculated EESC reaches the 1980 level again in about

2044. This is a few years earlier than estimated in

Madronich and Velders et al. (1999), 4 years longer than

if production of all ODSs were to cease in 2003, and 10

years longer than if emission of all ODSs were negligible

after 2003. In the maximum-emissions scenario (Am),

EESC drops to the 1980 level in about 2048.

The cases provide a gauge of maximum potential

reductions in EESC that would be achievable from further

restrictions to future production or future production and

emission of ODSs. Implementation of any new control

measures, however, must include a consideration of costs

and feasibility in addition to any atmospheric benefits, but

those considerations are not discussed here. From the zeroproduction cases presented in Table 1-17, the influence of

future ODS production allowances is apparent.

Allowances for continued production of HCFCs, methyl

bromide, and CFCs have the largest influence on integrated

EESC in the future. The zero-emission cases quantify the

contribution of future emissions plus future production to

future EESC (relative to the Ab scenario) and suggest that

ODS emissions from reserves of unemitted material will

affect EESC as much as or more than the current future

ODS production allowances. The exception is methyl bromide, where reserves are presumed to be negligible; its

influence on future EESC is primarily dependent upon the

magnitude of future production.

To show the effect of noncompliance with the

Montreal Protocol, the Pc scenario was constructed. This

scenario demonstrates that with continued, constant production of ODSs, even at the reduced amounts reported

for 1999, EESC remains well above the 1980 level through

2100 (Figure 1-23). In the Pc scenario, the integrated

EESC through 2100 is about 3 times as high as in the baseline scenario (Table 1-17). The results from this scenario

demonstrate clearly that atmospheric halogen will decline

to pre-Antarctic-ozone-hole conditions only with adherence to future limits on production of ozone-depleting

substances.

1.8.4 Uncertainty in the Baseline Scenario

The Ab scenario is constructed as a best guess for

future ODSs emissions and mixing ratios. There are several sources of uncertainties in these quantities.

First, uncertainty in the Ab scenario arises because

of our inability to predict future global HCFC production.

Limits apply to production or consumption of HCFCs in

countries operating under Article A5(1) of the Protocol

only after 2016 based upon amounts produced or consumed in 2015.

Second, emissions of carbon tetrachloride not

related to industrial CFC production are poorly characterSOURCE GASES

1.70

1940 1980 2020 2060 2100

0

500

1000

1500

2000

2500

3000

3500

EESC(ppt)

Ab Baseline

E0 Zero emissions

Am Maximum production

P0 Zero production

Pc Continued production

1980 level

E0

Am

Pc

Figure 1-23. Estimates of

future EESC based on the baseline scenario (Ab, solid line), the

maximum scenario (Am, longdashed line), and the hypothetical cases of zero emissions in

2003 and thereafter (E0) and

zero production in 2003 and

thereafter (P0) of all anthropogenic ODSs. Also shown are

results from the scenario with

continued ODS production in

the future at 1999 rates (Pc),

production that is substantially

larger than allowed in the fully

revised and amended Montreal

Protocol. Additional description

of these scenarios is found in

Tables 1-13 and 1-14.
ized. Accordingly, it is difficult to know accurately how

emissions not related to CFC production will change in

the future.

Third, there are uncertainties in the release rates of

ODSs to the atmosphere. In these scenarios, future annual

emissions are calculated as a fraction of the unemitted

reserves (or “bank”). This release fraction is based on an

analysis of the quantities present in, and emitted from, the

banks over the past decade (1990-1999). These release

fractions are held constant from 2000 onward.

Destruction of reserves or changes in the use patterns of

ODSs would change the estimated release fraction and

thereby the emissions and shape of the effective equivalent stratospheric chlorine (EESC) curve in the future, but

less so for the integrated EESC values. For integrated

EESC, when the emissions occur is less important; if the

ODS reaches the stratosphere before the EESC value falls

below the 1980 value (approximately 2040), it contributes

to the integrated value.

Fourth, the size of reserves of produced but unemitted ODSs is uncertain. In creating scenario Ab, historical emissions were adjusted so that the calculated mixing

ratios were consistent with measurements. These adjustments depend on the lifetimes and release rates to the

atmosphere of ODSs, and they affect our estimates of

present-day banks of unemitted ODSs. This is especially

true for Halon-1211, considering the large uncertainty in

reserves and in its global lifetime. To understand the influence these uncertainties have on the calculation of future

EESC, selected sensitivity tests were performed and compared with the results of the Ab scenario. If the lifetime of

Halon-1211 were 11 years, EESC would drop to the 1980

amount (integrated from 2002 onward) 2.0 years earlier,

and the integrated EESC would be 7.9% smaller than compared with the Ab scenario in which a 16-yr lifetime for

Halon-1211 is presumed. An additional future production

for this halon totaling 22.9 Gg over the years 2000 to 2009

(from Fraser et al., 1999) and a 35-Gg larger bank in 2000

yields an increase in integrated EESC of 3.5% (integrated

from 2002) and a delay in recovery of only 0.5 years.

Smaller Halon-1211 reserves in 2000 of 65 Gg (Fraser et

al., 1999; instead of 98 Gg in Ab) result in integrated EESC

being –3.4% compared with the Ab scenario.

Fifth, a simple box model is used for the calculation of mixing ratios in scenarios, as has been done in past

ozone Assessments. Advantages of the simple model are

that it maintains consistency with past reports and is more

easily interpreted than the 2-D and 3-D interactive models.

Simple box models have many limitations that affect the

accuracy of atmospheric mixing ratios calculated from

emissions. For example, the exponential decay time constant of a trace gas in the absence of emissions is the same

as the trace gas lifetime in a one-box model. In the real

atmosphere, however, this time constant can be different

from the steady-state trace gas lifetime (Prather, 1996).

Although such effects are thought to be small for longlived ODSs, they are predicted to be substantial for some

short-lived gases such as methyl bromide (Prather, 1997).

Sixth, accurate predictions of future halocarbon

mixing ratios rely on an understanding of loss rates in the

future (see Section 1.8.5). The composition of the atmosphere is changing, and this may alter the chemical or

dynamical processes that regulate lifetimes of ODSs (Krol

et al., 1998; Prinn et al., 2001; Butchart and Scaife, 2001).

1.8.5 Potential Influence of Future

Climate Change on Halogenated

Source Gases

About 80% of the equivalent chlorine in today’s

atmosphere arises from direct anthropogenic release of

ODSs. Climate change could influence future atmospheric amounts of these gases by affecting the rates at

which these gases are removed from the atmosphere

(Prinn et al., 2001; Butchart and Scaife, 2001). For CFCs,

halons, and CCl4, rates of atmospheric removal are governed by the flux of high-energy light in the stratosphere

and the rate at which air is mixed through the stratosphere.

If changes in climate were to affect rates of air transport

through the stratosphere, trace gas lifetimes would be

altered. The lack of predictive understanding of such

changes does not allow a reliable estimate of whether lifetimes would become longer or shorter in the future in an

atmosphere with elevated greenhouse gas abundances.

For methyl chloroform, HCFCs, methyl halides, and other

chemically reduced gases, climate change could influence

loss rate through changes in mean global concentrations

of the hydroxyl radical. This oxidant plays a key role in

determining the lifetime of these important ozonedepleting substances. Amounts of this short-lived, but

powerful atmospheric oxidant depend directly on

humidity, sunlight, temperature, and the abundance of

many different trace gases likely to be affected by climate

change and human behavior in the future.

The remaining 20% of ODSs in today’s atmosphere

arises from CH3Cl, CH3Br, and short-lived brominated

gases (see Section 1.2.4 and Chapter 2). These gases are

unique among ODSs because the predominant sources are

believed to be nonindustrial, and climate change could

affect the magnitude of sources for these gases. As a result

there is less direct control over mixing ratios of these gases

in the atmosphere. Our current understanding is limited

for predicting even the net change in sources and sinks

one might expect in the future for these methyl halides as

SOURCE GASES

1.71
a result of climate change or wide-scale land-use change.

The influence of temperature on air-sea concentration differences have been identified recently (Groszko and

Moore, 1998; King et al., 2000), but the relevance of this

information to trace gas fluxes in and out of the future

ocean has yet to be demonstrated.

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1.83
CHAPTER 2

Very Short-Lived Halogen and Sulfur Substances

Lead Authors:

M.K.W. Ko

G. Poulet

Coauthors:

D.R. Blake

O. Boucher

J.H. Burkholder

M. Chin

R.A. Cox

C. George

H.-F. Graf

J.R. Holton

D.J. Jacob

K.S. Law

M.G. Lawrence

P.M. Midgley

P.W. Seakins

D.E. Shallcross

S.E. Strahan

D.J. Wuebbles

Y. Yokouchi

Contributors:

N.J. Blake

J.H. Butler

A.R. Douglass

V.L. Dvortsov

I. Folkins

P.H. Haynes

A. Mellouki

M.J. Prather

J.M. Rodríguez

S.M. Schauffler

T.G. Shepherd

C. Textor

C. Timmreck

D.K. Weisenstein

CHAPTER 2

VERY SHORT-LIVED HALOGEN AND SULFUR SUBSTANCES

Contents

SCIENTIFIC SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1

2.1 INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.5

2.2 DYNAMICAL AND CHEMICAL CHARACTERISTICS OF THE UPPER TROPOSPHERE (UT)

AND THE TROPICAL TROPOPAUSE LAYER (TTL) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.7

2.2.1 Moist Convection and the Tropical Tropopause Layer . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.8

2.2.2 Stratosphere-Troposphere Exchange . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.12

2.2.2.1 Exchange Across the Tropical Tropopause . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.12

2.2.2.2 Exchange Across the Extratropical Tropopause . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.12

2.2.3 Chemistry in the UT and TTL . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.13

2.2.4 Chemical Tracer Observations for Model Evaluation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.14

2.3 ATMOSPHERIC CHEMISTRY OF HALOGENATED VERY SHORT-LIVED (VSL) SUBSTANCES . . . 2.15

2.3.1 Removal of the Halogen Source Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.16

2.3.2 Production and Removal of Degradation Products . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.18

2.3.2.1 Degradation Mechanisms for Source Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.18

2.3.2.2 Intermediate Products . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.19

2.3.2.3 Inorganic Bromine and Iodine Species . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.20

2.3.3 Degradation Mechanism for n-Propyl Bromide (n-PB) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.22

2.3.4 Stratospheric Iodine Chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.24

2.4 CONTRIBUTION OF HALOGENATED VSL SUBSTANCES TO THE STRATOSPHERIC

INORGANIC HALOGEN BUDGET . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.25

2.4.1 Observed Concentrations and Sources of Organic Chlorine, Bromine, and Iodine Compounds . . . 2.26

2.4.1.1 Chlorine Compounds . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.26

2.4.1.2 Bromine Compounds . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.29

2.4.1.3 Iodine Compounds . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.29

2.4.2 Estimates of Contributions from VSL Halogen Source Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.30

2.4.3 Modeling Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.32

2.4.4 Information Needed for Future Evaluation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.33

2.5 ESTIMATES FOR THE POTENTIAL IMPACT OF HALOGENATED VSL SUBSTANCES ON

COLUMN OZONE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.34

2.5.1 Review of Methodolgy for Calculating ODPs for Long-Lived Source Gases . . . . . . . . . . . . . . . . . 2.35

2.5.2 Modified Approach for Estimating ODP for Halogenated VSL Substances . . . . . . . . . . . . . . . . . . . 2.36

2.5.3 Existing Modeling Studies on ODP for n-PB . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.37

2.5.4 Information Required for Evaluating Ozone Impact from VSL Source Gases . . . . . . . . . . . . . . . . . 2.38

2.6 VSL SULFUR SPECIES AND STRATOSPHERIC AEROSOLS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.39

2.6.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.39

2.6.2 Chemistry of the VSL Sulfur Species . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.40

2.6.3 Sources for the Stratospheric Sulfate Layer . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.42

2.6.3.1 Direct Input from Explosive Volcanic Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.42

2.6.3.2 VSL Precursors Emitted in the Troposphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.42

2.6.4 How Can We Predict Future Changes in the Stratospheric Sulfate Layer? . . . . . . . . . . . . . . . . . . . . 2.43

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.44

SCIENTIFIC SUMMARY

Definition of Very Short-Lived Substances

Very short-lived (VSL) substances are substances that have chemical lifetimes comparable with tropospheric transport

time scales, with the result that the steady-state mixing ratio of the substance in the troposphere depends on where and when

(time of the year) it is released. In practice, this happens for species with atmospheric lifetimes of a few months or less.

Ozone Depletion Potential (ODP) for Halogen Source Gases

• The Ozone Depletion Potential (ODP) of a halogen source gas is defined as the time-integrated

effect on stratospheric ozone caused by emission of X at location and time te relative to the time-integrated effect

on stratospheric ozone from the same mass emission of CFC-11 (CCl3F) at the same location and time.

• In the general case, consists of two terms: and . 

(source gas injection (SGI) term) and (product gas injection (PGI) term) represent the effects

caused by inorganic halogen atoms that are transported to the stratosphere in the form of the source gas and in the

form of degradation products, respectively.

• For long-lived source gases, the value for is independent of and te, and is

negligible in comparison to . Thus, a single value of ODP can be obtained based on the atmospheric lifetime of the source gas independent of location and time of emission. These are the traditional singlevalue ODPs reported in the literature.

• For VSL source gases, the value depends on the location and time of emission. The

value depends also on the properties of the degradation products.

• If degradation of the VSL source gases gives rise to significant inorganic halogen concentration in the upper troposphere (UT), this may lead to ozone depletion in that region. The question then arises if this depletion should be

added to the ODP value, which traditionally refers only to depletion in the stratosphere. It also points to the need to

evaluate the role of halogen chemistry in the UT.

Transport from the Boundary Layer to the Stratosphere

• The most efficient route for irreversible transport of VSL source gases and their degradation products from the surface to the stratosphere is in the tropics, because the vertical transport times from the surface to the upper troposphere

are short, and air that enters the stratosphere through the tropical tropopause remains there for more than a year.

• In the tropics, the transition from tropospheric to stratospheric air takes place in a “tropical tropopause layer” (TTL)

between the secondary tropical tropopause (STT, 11-13 km) and the thermal tropical tropopause (TTT, 16-17 km).

Transport of VSL source gases and degradation products from the surface to the UT and the lower part of the TTL is

primarily through convection. From the TTL, a small fraction of the air enters the stratosphere via slow ascent

through the tropical tropopause or via isentropic transport into the extratropical lower stratosphere.

• Current estimates indicate that air at the base of the TTL is replaced by convection from the tropical boundary layer

on a time scale of 10-30 days. Thus, a significant fraction of the emitted VSL source gases can be expected to reach

the TTL in regions of deep convection.

• In the extratropics, exchange across the tropopause maintains a “transition layer” of 1-2 km with chemical characteristics intermediate between those of the UT and LS (lower stratosphere). Frontal lifting can transport air on time

scales of a few days to the UT or even occasionally into the LS. Through such mechanisms, VSL source gases are

expected to enter the transition layer.

ODPPGI(, ,) Xx t e e

r ODPSGI(, ,) Xx t e e

r

ODPSGI(, ,) Xx t e e

r ODPPGI(, ,) Xx t e e

r r x ODPSGI e (, ,) Xx t e e

r

ODPPGI(, , Xx t ) e e

r ODPSGI(, ,) Xx t e e

r ODPPGI(, ,) Xx t e e

r ODPSGI(, ) Xxt e e

r ODP( , , ) Xx t e e

r

r

xe

X Xx t e e ( ( , , )) ODP r

VERY SHORT-LIVED SUBSTANCES

2.1
Chemistry for VSL Halogens

• Inorganic chlorine, bromine, and iodine have the potential to destroy stratospheric ozone. For the same concentration in the stratosphere, iodine is most efficient in removing ozone, followed by bromine and chlorine. Laboratory

data on iodine stratospheric chemistry have led to downward revision of the iodine efficiency in depleting ozone in

the stratosphere. The estimated efficiency factor relative to chlorine (~150-300) is still higher than that of bromine

(~45).

• The main uncertainties in estimating the impact of VSL halogen source gases on stratospheric ozone lie in the physical and dynamical processes transporting halogen into the stratosphere and the behavior of their intermediate

degradation products; uncertainties in the gas-phase chemistry of the source gases are minor.

• The products expected from many of the bromine and iodine VSL source gases considered in this chapter are all

rapidly converted to inorganic halogen. The inorganic halogen multiphase chemistry for bromine and iodine is not

very well known. Heterogeneous processes that recycle reservoir species to reactive inorganic bromine and iodine

may significantly increase the efficiency of halogen injection (via the PGI pathway) into the stratosphere.

• The odd-hydrogen (HOx) chemistry in the upper troposphere is very critical both for the destruction/production of

ozone and for the delivery of VSL source gases to the stratosphere. The understanding of this chemistry requires a

better knowledge of the sources and processes for the oxygenated organics (carbonyl compounds, peroxides, etc.)

and for the nitrogen oxides (NOx) and total reactive nitrogen (NOy) species.

VSL Halogen Source Gases in the Present-Day Atmosphere

• The observed concentrations of VSL halogen source gases are typically a few parts per trillion (ppt) for bromine

and iodine VSL source gases, and up to 20 ppt for VSL chlorine source gases. The majority of these gases are of

natural origin.

• Two model studies simulated the atmospheric distribution of bromoform (CHBr3) assuming a simplified uniform

(space and time) ocean source. The calculated atmospheric distributions indicate that an average surface mixing

ratio of 1.5 ppt at the surface could maintain about 1 ppt of inorganic bromine in the stratosphere. The simulation

shows that one-half to three-fourths of the bromine enters the stratosphere in the form of inorganic products. These

results can be used as a guideline for estimating contributions from other source gases that have local photochemical lifetimes of about 10 days, no long-lived intermediate degradation product, and relatively uniform sources over

the globe.

• Very short-lived bromine and iodine source gases with surface concentrations of a few ppt could have large effects

on the current inorganic bromine and odd-iodine budget because the stratospheric concentrations of inorganic

bromine and iodine are about 20 ppt and less than 1 ppt, respectively. The transfer of inorganic bromine from the

UT may account for the current discrepancy in the stratospheric inorganic bromine budget (as discussed in this

chapter and in Chapters 3 and 4).

Modeling Studies to Determine Values for the ODP of n-Propyl Bromide

• n-propyl bromide (n-PB; CH3CH2CH2Br) is removed by reaction with OH, with local photochemical lifetimes in

the tropical troposphere of about 10-20 days. Laboratory data indicate that its degradation products, and particularly bromoacetone, have lifetimes shorter than a day.

• Given the complexity of atmospheric processes, three-dimensional numerical models are the preferred tool to evaluate the ODP for VSL source gases. Results for n-PB from three separate model studies that computed

were assessed. These results can be used as a guideline for other gases with similar local photochemical lifetimes (~10 days) and degradation schemes.

ODP n PB ( , ,) - rx t e e

VERY SHORT-LIVED SUBSTANCES

2.2
• Two of the three studies provided values only for . In the third study, which computed contributions from , the values of are 0.04 for uniform emission over land. In this

case, 0.5% of the bromine emitted as n-PB reaches the stratosphere, with about two-thirds entering in the form of

degradation products. The model simulations also give 0.1 for tropical emission and 0.02-0.03 for emissions

restricted to Northern midlatitudes.

VSL Sulfur Species and Stratospheric Sulfate

• A sulfur (S) source of 0.06-0.15 Tg (S) yr-1 is needed to maintain the observed burden of the background stratospheric sulfate layer during periods with low volcanic activity. Current model estimates indicate that degradation

of carbonyl sulfide (OCS) in the stratosphere can account for 20-50% of the needed source, with the balance from

transport of sulfur VSL gases (dimethyl sulfide (DMS; CH3SCH3), hydrogen sulfide (H2S), sulfur dioxide (SO2))

and sulfate formed in the troposphere.

• Our current best estimates for the surface sources of sulfur are 10-30 Tg (S) yr-1 for DMS, 1-2 Tg (S) yr-1 for OCS,

0.5-1.5 Tg (S) yr-1 for CS2, 6-9 Tg (S) yr-1 for H2S, and 67-100 Tg (S) yr-1 for SO2. Emission from aircraft engines

provides a source of 0.06 Tg (S) yr-1 in the upper troposphere and lower stratosphere.

• Model simulations show that a few tenths of one percent of the SO2 emitted at the ground reaches the stratosphere.

This is consistent with estimates for VSL halogen source gases. For DMS, multiphase oxidation may occur and

reduce its possible transfer to the UT.

• Previous climate calculations assumed that the change in stratospheric sulfate loading is proportional to changes in

total SO2 emission. Changes in the spatial distribution of the emission can have a large effect. Depending on how

much of the sulfur gas is converted to small sulfate particles in the UT, the change in surface area per unit mass of

sulfur transferred to the stratosphere could differ.

Model Evaluations

• In order to evaluate the contribution of VSL source gases to halogen and sulfur loading, global 3-D model simulations are required. Simulating aerosol microphysics in the troposphere with very short time constants in globalscale models presents a large challenge. Significant uncertainties exist in the treatment of dynamical and physical

loss processes in such models.

• Further evaluation of model performance against observations is necessary, particularly relating to rapid transport

processes and treatments of cloud scavenging. Various tracers can be used to test specific components: radon-222

and short-lived hydrocarbons for continental convection; methyl iodide, dimethylsulfide, and ozone for marine convection; lead and beryllium isotopes, and soluble gases for precipitation scavenging; and ozone for stratospheretroposphere exchange.

ODP n PB ( , ,) - rx t ODP n PB PGI e e ( , ,) - r

x t e e ODP n PB SGI( , ,) - r

x t e e

VERY SHORT-LIVED SUBSTANCES

2.3

VERY SHORT-LIVED SUBSTANCES

2.5

2.1 INTRODUCTION

An important component of previous Assessments

(e.g., WMO, 1999, 1995) is the discussion of the observed

concentrations, budgets, trends, and atmospheric lifetimes

of source gases that affect stratospheric ozone. These

gases are mostly emitted at the surface. The source gas

and its degradation products are transported to the stratosphere, where they could affect the ozone abundance.

Examples of these include nitrogen species from nitrous

oxide (N2O); chlorine, bromine, and iodine species from

organic halogen-containing chemicals; and methane

(CH4) reacting to produce water vapor in the stratosphere.

Sulfur species found in the troposphere (e.g., carbonyl

sulfide (OCS), carbon disulfide (CS2), hydrogen sulfide

(H2S), dimethyl sulfide (DMS, CH3SCH3), sulfur dioxide

(SO2)) provide the source of sulfur to maintain the stratospheric sulfate layer. The sulfate aerosol particles provide

sites for heterogeneous reactions that control the partitioning of the nitrogen and halogen radicals. One aim of

the Assessments is to establish the tools that enable one to

predict how changes in emissions of the various source

gases would affect ozone. In the case of the halogen

species, it is done in terms of the equivalent chlorine concentration estimated from emissions and lifetimes. A similar connection could be made for the sulfur source gases

if one could predict how the stratospheric sulfate may

change in response to changes in emissions. In this

chapter, we restrict the discussion to halocarbons and the

sulfur gases. Other species with similar lifetimes are

affected by many of the same processes discussed in this

chapter. One example is oxides of nitrogen released at

the ground (from combustion and biomass burning) and

produced in the free troposphere (by lightning or aircraft

engines). They are not included in this chapter.

The purpose of this chapter is to discuss how very

short-lived (VSL, defined later in this section) substances

released at the surface or produced in situ in the troposphere

could affect ozone in the stratosphere. Because the subject

is relatively new, this chapter will deviate from the standard

approach of providing a detailed review and assessment of

the information in the published literature. The current state

of our understanding does not provide concrete evaluations

to assess the merit of the approaches and make recommendations for specific values, e.g., of Ozone Depletion

Potentials (ODPs). Instead, the chapter focuses on outlining

viable approaches with emphasis on how observations may

be used in future evaluations. If certain simplifying assumptions are made, order-of-magnitude estimates for some of

the parameters can be provided so that policymakers can

start to consider how best to approach the problem. The

emphasis of the chapter is on the issues needed to be

resolved and the observations needed for evaluation.

In the discussion, it is often convenient to distinguish among the different forms in which the halogen

atoms and sulfur atoms are transported to the stratosphere.

Halogen atoms bound in the source gases and the intermediate degradation products are referred to here as

organic halogen. The final products are referred to as inorganic halogen; these include, in the case of chlorine, the

reservoir species (hydrogen chloride (HCl), chlorine

nitrate (ClONO2), hypochlorous acid (HOCl)) and the radical species (atomic chlorine (Cl), chlorine monoxide

(ClO)). For sulfur, an analogous approach would identify

OCS, CS2, H2S, and DMS as source gases, and SO2 and

various forms of sulfate (SO4) in gaseous, liquid, or solid

form as the final product. As seen in Section 2.6, this turns

out not to be desirable, because SO2 and SO4 are also

released at the surface. We will refer to a sulfur species

as sulfate or sulfate precursor in the discussion.

A major goal in the study of halogen source gases

has been to track the life cycle of a source gas so that one

can evaluate how its emission may affect the inorganic

halogen budget in the present-day stratosphere. This is

accomplished by combining the amount of the source gas

emitted at the surface with the percentage of the emitted

halogen atoms that is delivered to the stratosphere. It is

important to provide values for the latter because it allows

one to define the relative ODP of an individual source gas.

Traditional methods for evaluating the contribution of

organic halogen source gases with lifetimes of 1 year or

longer have been reviewed and assessed in previous

Assessments (e.g., WMO, 1995, 1999). However, such

methods cannot be applied to VSL species. To understand

why, we examine the processes that deliver the final products to the stratosphere. For ease of discussion, we will

present the arguments in the context of the halogen

species. It should be noted that many of the ideas are

equally applicable to the sulfur species.

Figure 2-1 illustrates that there are two pathways

for delivering the final products to the stratosphere. For

the source gas injection (SGI) pathway, the source gas is

transported to the stratosphere and then reacts to release

the halogen atoms. For a given emission location and

time, its efficiency depends mostly on the properties of

the source gas (essentially its removal rate by chemical or

physical processes in the troposphere as compared with

the transport rate from the ground to the stratosphere).

The product gas injection (PGI) pathway involves transport to the stratosphere of intermediate or final products

produced in the troposphere. Its efficiency depends also

on the properties of the degradation products.

To put things in perspective, we provide some

rough numerical estimates for the two pathways for longlived source gases that have uniform mixing ratios in the
troposphere. Long-lived source gases that are inert in the

troposphere (most chlorofluorocarbons) are recycled

between the troposphere and stratosphere until they are

all removed via photochemical reactions in the stratosphere. Thus, 100% of the final product is delivered to

the stratosphere via the SGI pathway. For long-lived

source gases that are removed efficiently in the troposphere (most hydrogenated chloroflurocarbons), at least

90% of the mass burden is found in the troposphere at

steady state. If we assume that the local lifetimes in the

troposphere and stratosphere are similar, this would imply

that 90% of the final product is released in the troposphere.

Thus, only 10% is delivered to the stratosphere via the

SGI pathway. If one assumes that less than 1% of the

degradation products (in the form of intermediate degradation products or inorganic halogen) released in the troposphere is transported to the stratosphere, then the supply

from the PGI pathway (0.009 of emission) is a factor of

10 smaller than from the SGI pathway (0.1 of emission).

This is why the contribution from the PGI pathway is usually ignored in considering the long-lived source gases.

We now explain the criteria for classifying a substance as VSL, as distinct from those discussed in Chapter

1 of this report. The established method for estimating

inorganic halogen loading from the SGI pathway depends

on the observed fact that the mixing ratio of the long-lived

source gas is essentially uniform in the troposphere. [Note

that the key issue is the difference in mixing ratio between

the boundary layer (BL) and the upper troposphere (UT).

Asmall interhemispheric difference (~20%) in the BL will

not affect the accuracy.] In these cases, the fraction of the

emitted halogen atoms entering the stratosphere is independent of the location and time of emission. The inorganic halogen loading in the stratosphere can be estimated

from the steady-state burden of the source gas, which is

related to its atmospheric lifetime. Details of the transport processes in the troposphere play no role. As the local

lifetime of the source gas gets shorter, its mixing ratio will

no longer be uniform in the troposphere. The method used

for long-lived species is no longer accurate when the

mixing ratio in the UT is less than half of the mixing ratio

in the BL. A survey of chemicals in the atmosphere indicates that this occurs when the local lifetime (inverse of

the local removal rate) is 0.5 years or shorter. This will

serves as the criteria for identifying VSL species.

The fraction of the emitted halogen atoms that are

transported to the stratosphere via the PGI pathway

depends on the properties of the intermediate and final

products as well. Consider first the situation where the

intermediate products are so short lived that, for practical

purposes, the inorganic halogen species can be treated as

being produced by the initial reaction experienced by the

source gas. Many of the inorganic halogen species are

water soluble or will react in solution to form watersoluble products. They are removed efficiently in the lower

troposphere, with a local residence time of about 10 days.

VERY SHORT-LIVED SUBSTANCES

2.6

Intermediate

degradation

product

Intermediate

degradation

product

Final

product

Final

product

Emission

Source

gas

Source gas

Stratosphere

Troposphere

Chemical transformation Washout Strat/trop exchange Convection SGI Pathway PGI Pathway PGI Pathway

Figure 2-1. Schematic diagram showing the two pathways through which a source

gas released at the surface

delivers its degradation products to the stratosphere. The

orange arrows and green

arrows represent transport

into and out of the stratosphere, respectively. The two

pathways are labeled SGI for

source gas injection and PGI

for product gas injection.
In order for the PGI pathway to provide an effective source

to the stratosphere, the products should be produced at least

as fast as the washout rate (Ko et al., 1997). Thus, the PGI

pathway will be important if the local photochemical lifetime of the source gas is of the order of 10 days or shorter.

The properties of the intermediate products could

change the estimate in a significant way. If the intermediate products are long lived in the troposphere (i.e., with

a photochemical lifetime longer than the source gas or the

washout lifetime of the inorganic halogen), they could

serve as a vehicle for transporting halogen atoms to the

stratosphere and deliver much more of the degradation

products via the PGI pathway. An example of an intermediate product having lifetime longer than the source

gas is phosgene (COCl2) formed from degradation of

trichloroethene (C2HCl3). At the same time, if the local

photochemical lifetimes of the intermediate products in

the stratosphere are longer than their residence times in

the stratosphere, they may be removed from the stratosphere before they release their halogen atoms. This will

effectively decrease the efficiency of the SGI pathway

(see Kindler et al., 1995).

Many of the same issues apply to the VSL sulfate

precursor gases and how their fates in the troposphere

would affect the stratospheric aerosol layer. In addition

to the release of source gases at the surface, there is a direct

way of injecting sulfur species into the stratosphere

through explosive volcanic eruptions. In this chapter, we

will briefly review the effects of volcanoes, building on

work from the previous Assessment (Godin and Poole et

al., 1999). The major emphasis here is on the role of the

VSL precursor gases.

The structure of this chapter is as follows. Section

2.2 discusses the transport mechanisms in the troposphere

and cross-tropopause exchange, and the chemical environment in the upper troposphere. Section 2.3 describes

the removal mechanisms of the VSL halogen source gases

and their degradation products. It also includes a brief

discussion of the ozone removal efficiency of iodine in

the stratosphere. Section 2.4 discusses how VSL halogen

source gases may contribute to the inorganic halogen

budget in the current stratosphere. The observations for

these source gases are reviewed, with emphasis on our

understanding of the budgets. Section 2.5 focuses on the

issue of evaluating the inorganic halogen loading estimates and ozone depletion estimates of VSL halogen

source gases. In addition to natural species, the discussion includes industrial species and the information that

is needed on emission inventories to estimate their

impact. Finally, Section 2.6 discusses the effects of sulfurcontaining VSL gases and how they affect future trends in

the stratospheric sulfate layer.

2.2 DYNAMICAL AND CHEMICAL CHARACTERISTICS OF THE UPPER TROPOSPHERE (UT) AND THE TROPICAL

TROPOPAUSE LAYER (TTL)

The critical parameters for understanding the

processes that control fluxes of key trace gases from the

troposphere to the stratosphere are the mixing ratios of

the trace gases in the upper troposphere (UT) and in the

tropical tropopause layer (TTL) and the transport rates

across the tropopause. We employ the following definitions in this chapter: the UT is the region from about 8-12

km, which is capped by the tropopause in the extratropics,

and which overlaps with the lowest part (~1 km) of the

TTL in the tropics. The TTL is bounded above by the

thermal tropical tropopause (TTT; defined here as the cold

point of the sounding) at about 16-17 km, and bounded

below by a secondary tropical tropopause (STT) at 11-13

km. The STT is related to large-scale circulation in that

it corresponds to the maximum Hadley-cell outflow

(Highwood and Hoskins, 1998). At these levels, convection is no longer able to maintain a moist adiabatic temperature lapse rate, owing to the rapid decreases with

height of the cumulus mass flux, and IR radiative cooling.

Within the TTL there is a gradual transition from tropospheric to stratospheric dynamical and chemical characteristics. However, the rapid decrease with height of the

convective outflow, and the rapid transition from clear-sky

radiative cooling to radiative heating, combine to make

the transition at the TTT fairly sharp (see, e.g., Folkins et

al., 1999; Hartmann et al., 2001; Jensen et al., 2001;

Thuburn and Craig, 2000). The TTL is a relatively new

concept, and its nature and regional variations are still

under investigation. The concentrations of trace gases in

the UT and TTL are controlled by transport, physical and

chemical removal processes in the troposphere and lower

stratosphere, and the location and timing of emissions.

We consider two broad classes of transport mechanisms, illustrated in Figure 2-2. The first is deep moist

convection, which transports air rapidly from the

boundary layer to the UT, with a fraction detraining well

into the TTL and some directly into the stratosphere. The

second is large-scale stratosphere-troposphere exchange

(STE), which transports air from the TTL upward into the

stratospheric overworld, and laterally into the extratropical lower stratosphere. In addition, this section includes

a discussion of extratropical STE, the chemical environment in the UT and TTL, and a survey of the database of

observations that are needed to critically evaluate the

results of models.

The numerical simulation of very short-lived substances is difficult because of the wide range of scales on

VERY SHORT-LIVED SUBSTANCES

2.7
which relevant processes occur, ranging from microphysical (uptake of gases in cloud droplets) to synoptic and

global (advection). Global models are hindered by their

coarse spatial and temporal resolutions (which are typically at best of the order of 100 km in the horizontal, 0.5

km in the vertical in the free troposphere and 0.1 km in

the BL, and 10 minutes in time). This results in the need

for “parameterizations” of important subgrid-scale

processes. Regional models are able to resolve smallscale processes better, but their limited spatial domain

requires the specification of boundary conditions, which

makes it difficult to extrapolate to global-scale estimates.

In one comparison of simulated results from global tropospheric chemistry transport models with ozone (O3)

data from the Measurement of Ozone and Water Vapor by

Airbus In-service Aircraft (MOZAIC) program, it was

found that the models generally smeared out observed gradients, that they performed less well in regions where

meteorological analyses, used as input, are expected to be

less reliable, and that higher resolution seemed more

important than an extensive hydrocarbon chemistry

scheme for the quality of the simulated upper tropospheric

O3 (Law et al., 2000).

2.2.1 Moist Convection and the Tropical

Tropopause Layer

Moist convection is characterized by rapid updrafts

(driven by latent heat release from condensing water) with

sustained velocities of up to several meters per second

over regions of the order of a kilometer in diameter. The

updrafts are balanced by downdrafts (largely driven by

evaporating precipitation) with a similar spatial extent,

and by large-scale subsidence over a much larger area

where clear-sky radiative cooling occurs. In the UT (particularly the tropics), where moist convective adjustment

holds, trace gases are strongly influenced by rapid upward

transport within convective clouds and slow descent outVERY SHORT-LIVED SUBSTANCES

2.8

5 km 

TTL

TTT

STT

WINTER SUMMER

VSLS

VSLS

VSLS VSLS IO, BrO

XHal

cirrus

XHal

I, Br,

Cl

IO, BrO,

ClO

ozone

loss

OH OH

hν hν

OH

hν

XHal

 1-2 km 

11 km 

20 km 

28 km 

PBL

EQUATOR

VSLS

washout

washout

XHal

XHal

hν

Key

VSLS - Very Short-Lived Source gases

and organic intermediates

XHal - intermediate products and

inorganic halogen reservoirs

XO - halogen radicals, X=Cl, Br, I

TTT - Thermal Tropical Tropopause

TTL - Tropical Tropopause Layer

STT - Secondary Tropical Tropopause

ExTL - Extra-tropical Transition Layer

PBL - Planetary Boundary Layer

XHal

Chemical and Transport Processes Affecting VSLS

convection subsidence

cross-tropopause

exchange

clouds

SO , DMS HSO OH SO 2 4

2-

x wet deposition dry/wet

deposition

Sulfur

chemistry

large-scale

transport

frontal

uplift

ExTL

OH

XO

IO, BrO IO, BrO

IO, BrO

Figure 2-2. Chemical and dynamical processes affecting VSL substances in the troposphere and stratosphere. Figure courtesy of K.S. Law, R.A. Cox, and P.H. Haynes, Cambridge University, U.K.
side convectively active regions, and nearly all air parcels

have had recent contact with the tropical boundary layer.

Convective clouds are often associated with the formation of precipitation (rain, snow, and graupel) and anvil

cirrus clouds. This leads to efficient scavenging of soluble gases by precipitation where convection is active.

Convective events are highly stochastic. An individual cloud can transport air parcels (thermals) from the

BL to the UT in less than an hour. However, the actual

concentration of the source gases and degradation products in the UT and TTL depends on how quickly they are

evacuated from the BL, how often thermals reach and penetrate the UT and the TTL, and how much old air from the

midtroposphere (which has much smaller concentrations

of the source gas) is entrained into the convective cores.

An estimate for the mean evacuation time for

the BL in the tropics can be obtained using relatively

well-defined values for the Hadley cell circulation and

the clear-sky radiative cooling. The observed annual

mean mass flux from the boundary layer in the Hadley

circulation between 15°N and 15°S is about 1011 kg s–1

(Hartmann, 1994). This, however, is the net mass flux

arising from the difference between the upward flux in

convection and the clear-sky subsidence. The clear-sky

radiative cooling rate in the tropical lower troposphere is

~1.6 K day–1 (e.g., Hartmann et al., 2001). Because convective disturbances cover less than 10% of the tropics,

the above cooling rate implies a downward mass flux

between 15°N and 15°S of about 2.0 ¥ 1011 kg s–1

(assuming descent along a dry adiabat). Thus the upward

flux from the boundary layer in convection must be about

3.0 ¥ 1011 kg s–1. Assuming that the BL has a 100-hPa

depth, then the mass between 15°N and 15°S is about

1.3 ¥ 1017 kg. The upward flux from the BL then implies

an evacuation time of about 5 days if one assumes that the

material is taken from the whole BL rather than just from

the top of the BL. On the basis of a 3-D general circulation model simulation, Prather and Jacob (1997) estimated

a mean evacuation rate of the 130-hPa-deep BL in the

tropics of 18% day-1 (see the GISS results in Figure 2-3),

which implies an e-folding removal time of 5.5 days, consistent with the above estimate. This estimate is at the

low end of the 5-10 day estimate made in Kley et al. (1996)

based on data from a different general circulation model,

and is slightly longer than the estimate of 4 days made by

Cotton et al. (1995). The actual evacuation rate will vary

regionally, depending on the strength and frequency of

convective events.

The GISS model results in Figure 2-3 show that the

magnitude of the mass flux across the STT (1.4 g m–2 s

–1)

is about half that of the mass flux out the BL (2.8 g m–2

s–1). As indicated above, the latter mass flux implies a

replacement time for the BL of about 5 days. Since the

TTL and BL have similar depths of about 100 hPa, the

average replacement time for the whole TTL by air

crossing the STT should be of the order of 10 days.

However, the replacement time is not uniform with altitude. Much of the detrainment in the UT and TTL occurs

near the STT, and within the TTL the mass flux decreases

rapidly with height, but the precise height dependence has

not yet been reliably determined. The local replacement

time (defined as the replacement time of a layer of air by

detrainment from convective clouds) can be estimated by

assuming that the upward mass flux from convection is

balanced by the clear-sky subsidence. Folkins et al. (2002)

used this method to obtain replacement times of 5 days

near the STT, and 100 days at 16 km. An alternative estimate can be obtained by using measured profiles of CO

and O3 (Dessler, 2002), yielding 10-30 days for the lower

TTL and 50-90 days for the upper TTL. It is difficult to

quantify the uncertainty in these estimates at present, but

it is likely that global models with different contemporary

convection parameterizations could produce numbers that

differ by at least a factor of 2.

The replacement time can be used to estimate how

much additional chemical transformation has occurred for

the VSL species in air parcels since they have left the BL.

The actual concentration of the species will depend also

on the entrainment of old air as the convective column

ascends the free troposphere. If the BL air has been diluted

by 50% during the ascent, the replacement time for the

TTL by BL air will be a factor of 2 larger than the estimates cited above. Cloud-resolving model simulations

(Lu et al., 2000) indicate that only a very small amount of

tracers that are produced in situ in the middle troposphere

are entrained into the updrafts, but instead are forced aside

and downward due to rising thermals. The wealth of literature on studies using cloud-resolving models (e.g., Lu

et al., 2000; Redelsperger et al., 2000; Bechtold et al.,

2000) reveals the immense complexity of transport in

cumulus clouds and the difficulties of representing this in

large-scale models.

The general transport characteristics of convection

have also been studied using idealized tracers in a 3-D

global chemistry-transport model, with a fixed uniform

mixing ratio at the surface and a specified local decay rate

(Crutzen and Lawrence, 2000). It was found that an insoluble gas with a lifetime of 30 days was 40-50% as abundant in the tropical UT as at the surface, whereas with a

lifetime of 10 days the tropical UT abundance was 25-

30% of the surface mixing ratio (see Figure 2-4). Lower

values were computed in the UT outside the tropics: 25-

35% and 10-15% for the 30- and 10-day lifetime tracers,

respectively. As discussed above, these estimates depend

VERY SHORT-LIVED SUBSTANCES

2.9
strongly on the respective model formulations, and thus

could differ significantly from the results obtained with

other models. Such modeling studies, unfortunately, are

rarely repeated once the first ones have been published.

The situation is even more complex for soluble

gases because of the variable relationship between updraft

mass flux and the amount of precipitation. Crutzen and

Lawrence (2000) also examined this issue using idealized

soluble tracers. Compared with the calculated concentration of an insoluble gas, they found that about 15%, 50%,

and 85% of gases with Henry’s Law solubility constants

of 103

, 104, and 105 M atm–1, respectively, were scavenged

by precipitation before they reached the UT (see Figure

2-5). This result was nearly independent of season, geographical location, and the decay rate applied to the

tracers, but was strongly dependent on the assumptions

made about whether tracers were rejected from, retained

in, or additionally scavenged by ice, snow, and graupel.

Similar studies using cloud-resolving models (Barth et

al., 2001; Yin et al., 2001) have also indicated that it is

important to consider kinetic limitations to trace gas

uptake in precipitation, because assuming Henry’s Law

equilibrium, as is generally done for the long time steps

of global models, will lead to an overestimate of the

amount of gases that can be taken up into precipitation,

and thus to the scavenging loss of highly soluble gases. A

promising approach for future studies is the use of observations of peroxides with various solubilities together with

insoluble tracers like CO to test parameterizations of convective scavenging (e.g., Mari et al., 2000). At present it

is difficult to apply these results to inorganic halogenated

species because of a lack of observations, combined with

the fact that they are also produced in situ in the UT and

can have highly variable mixing ratios at the ground.

VERY SHORT-LIVED SUBSTANCES

2.10

1000 0.9 950

870

390

255

100

GISS II

1.4

0.2

1.2

1.9wet convect. flux g m-2 s-1 flushing rate

17

 8

11 pressure (hPa) altitude (km)

Convective Turnover of Tropical Troposphere

1.8

0.8

Oslo/EC

 13 d 10 d

 5 d

 21 d 13 d wet convect. flux g m-2 s-1 flushing rate

STT

BL

Figure 2-3. Convective turnover rates in the tropical troposphere. The results are from two models, GISS II

(Goddard Institute for Space Studies), using winds generated by a general circulation model, and the Oslo

CTM, driven by European Centre for Medium-Range Weather Forecasts (ECMWF) winds; based on Prather

and Jacob (1997). The gray shaded outline provides a schematic representation of the structure of the convective cloud. The arrows represent entrainment fluxes (lower two arrows) and detrainment fluxes (top three

arrows). The average convective mass fluxes (in g m–2 s–1) and the implied turnover time constant (inverse of

turnover rate; in days) of the reservoirs are given for each model (in red for the GISS II model and in blue for

the Oslo model). The column mass densities in each reservoir are: 1.3 ¥ 106 g m–2 for the boundary layer

between 870 and 1000 hPa; 1.4 ¥ 106 g m–2 for the upper troposphere located between 390 and 255 hPa; and

1.6 ¥ 106 g m–2 for the tropical tropopause layer located between 255 and 100 hPa. The values from the GISS

II model are calculated for the region between 12°S and 12°N for May-August conditions, whereas those for

the Oslo model are for 22°S-22°N for January-December 1997. Figure courtesy of M. Prather, University of

California-Irvine, U.S.
Although this section has focused on the role of

deep convection, it is important to note that other

processes can also contribute significantly to the transport of air parcels from the BL into the mid- and upper

troposphere, especially in the extratropics. In particular,

recent attention has been given to forced uplift associated

with synoptic-scale frontal systems, which may be the primary mode of vertical transport in many regions of the

extratropics (Stohl, 2001; Zahn, 2001; Kowol-Santen et

al., 2001). Because frontal lifting is also associated with

clouds and precipitation, similar uncertainties for the scavenging of soluble gases as discussed above will apply.

The issue of frontal lifting is likely to see considerable

progress prior to the next Assessment. Other processes,

VERY SHORT-LIVED SUBSTANCES

2.11

τ = 30 d

τ = 10 d

0.0 0.2 0.4 0.6 0.8 1.0

100.0

1000.

200.

50.

500.

τ = 30 d

τ = 10 d

0.0 0.2 0.4 0.6 0.8 1.0

100.0

1000.

200.

50.

500.

Mass Mixing Ratio (ng/g)

Pressure (hPa)

EQ

Mass Mixing Ratio (ng/g)

Pressure (hPa)

45°N

Figure 2-4. Zonal mean vertical profiles of surface tracers at the equator and 45°N, plus the standard deviation around the respective latitude bands (horizontal bars), for insoluble tracers with decay lifetimes (τ) of

30 and 10 days; computed with the Model of Atmospheric Transport and Chemistry–Max-Planck-Institute for

Chemistry (MATCH-MPIC) for July 1993. Based on Crutzen and Lawrence (2000).

Hx = 0

Hx = 103

Hx = 104

Hx = 105

Hx = 106

0.0 0.2 0.4 0.6 0.8 1.0

100.0

1000.

200.

50.

500.

Hx = 0

Hx = 103

Hx = 104

Hx = 105

Hx = 106

0.0 0.2 0.4 0.6 0.8 1.0

100.0

1000.

200.

50.

500.

Mass Mixing Ratio (ng/g)

Pressure (hPa)

EQ

Mass Mixing Ratio (ng/g)

Pressure (hPa)

45°N

Figure 2-5. Zonal mean vertical profiles of surface tracers at the equator and 45°N for tracers with decay lifetimes of 30 days and Henry’s Law solubilities varying from Hx = 0 to Hx = 106 M atm–1 (insoluble to highly soluble); computed with MATCH-MPIC for July, 1993. Based on Crutzen and Lawrence (2000).
such as variations in the BL height in association with

non-zero large-scale vertical velocities, as well as topographical effects, can also contribute to the evacuation of

the BL.

2.2.2 Stratosphere-Troposphere Exchange

Most air arriving in the lower stratosphere passes

through the tropical tropopause (e.g., Holton et al., 1995),

although net transport of chemical constituents across the

subtropical and extratropical tropopauses is also known

to take place. Because the net large-scale motion is

upward in the tropical lower stratosphere and downward

in the extratropical lower stratosphere, air entering the

lower stratosphere via the tropical tropopause is believed

to have a larger residence time (>1 year) in the stratosphere. The precise mechanisms of the cross-tropopause

transport and their consequences for short-lived substances are still controversial. Furthermore, it is also

important to know the form of halogenated gases in the

UT, because soluble gases are subject to removal from the

UT by precipitation (as discussed in the previous section),

as well as by slower sedimentation of cirrus ice (Lawrence

and Crutzen, 1998), both of which can occur at rates that

can compete with the residence time of gases in the TTL.

In this section, we discuss how one may obtain estimates

of the flux across the tropopause by examining the mixing

ratios of the species near the tropopause and the mass flux

across the tropopause.

2.2.2.1 EXCHANGE ACROSS THE TROPICAL

TROPOPAUSE

The transition from air having compositions typical of the troposphere to air typical of the stratosphere

does not occur abruptly across the tropical tropopause

(TT). Above the level where the clear-sky mass flux is

mostly upward, air parcels are rising slowly into the stratosphere, driven by the Brewer-Dobson circulation (speed

of order 30-50 m day–1). However, some deep convection penetrates deep into the TTL, and a small fraction of

the air leaving the boundary layer (about 0.1%, based on

Gettelman et al., 2002) is directly lofted across the TT.

Thus, a small fraction of the air encountered directly above

and below the TT may have a distinct “BL character.” At

the same time, convective overshooting and breaking

waves can mix material back downward (Fujiwara et al.,

1998; Sherwood and Dessler, 2001), and therefore not all

the air in the TTL is transported irreversibly into the stratosphere. Notably, the radiative properties of the TTL are

sensitive to the stratospheric O3 distribution (Thuburn and

Craig, 2000), which is in turn sensitive to the transport of

constituents through the TTL.

It is also likely that some of the air in the TTL has

a stratospheric origin, entering via lateral mixing from the

extratropical stratosphere (Figure 2-2). A 20-yr climatology of intrusions from the extratropics into the tropics

(Waugh and Polvani, 2000) showed that Rossby wavebreaking events can result in transport of tongues of extratropical air deep into the tropics, and a more detailed look

at four Rossby wave-breaking events (Scott et al., 2001)

was used to “infer a potentially significant contribution to

the global stratosphere-troposphere exchange from the

wave-breaking processes.” Studies of mixing in the other

direction also help give an indication of the time scales

for two-way mixing, which can be comparable with convective time scales. For example, Dethof et al. (2000)

reported a contour advection study using European Centre

for Medium-Range Weather Forecasts (ECMWF) water

vapor fields, in which they found that the isentropic

transport of water vapor into the extratropical lower

stratosphere is similar in magnitude to upward transport

in tropical regions. Furthermore, the work of Strahan et

al. (1998) also showed evidence of such isentropic transport based on the seasonal cycle of carbon dioxide simulated in a global chemical transport model.

Finally, according to Rosenlof and Holton (1993), the

annual flux across the 100-hPa surface (i.e., TT) in the 15°N

and 15°S latitude range is 8.5 ¥ 109 kg s–1. Because this is

only ~3% of the flux of air out of the tropical BL computed

above (Section 2.2.1), the vast majority of the air parcels in

deep convection are circulated back down into the free troposphere to eventually re-enter the BL, or are transported

isentropically into the extratropical lower stratosphere. A

quantification of these fractions is not yet possible.

2.2.2.2 EXCHANGE ACROSS THE EXTRATROPICAL

TROPOPAUSE

Exchange of air across the extratropical tropopause

is accomplished by processes on a large range of temporal

and spatial scales. Although the net (downward) mass

exchange rates are controlled by global-scale processes as

part of the Brewer-Dobson circulation (Holton et al., 1995),

two-way exchange occurs on synoptic and sub-synoptic

scales and may have important consequences for chemical

species. Synoptic-scale frontal systems and to a lesser

extent midlatitude deep convection frequently bring constituents from the boundary layer to the upper troposphere,

leading to strong chemical contrasts across the tropopause.

These processes may also transport trace constituents,

including VSL substances, their degradation products and

inorganic halogen into the lowermost stratosphere.

Previous analysis of observed and modeled carbon

dioxide (CO2) levels (Bolin and Bischof, 1970; Strahan et

VERY SHORT-LIVED SUBSTANCES

2.12
al., 1998) suggested that there is not a large degree of communication between the UT and lower stratosphere (LS) in

midlatitudes because of a clear separation in the seasonal

cycle in these different regions. It was postulated that this

is due to the bulk of air in the lower extratropical stratosphere originating either through the tropical stratosphere

or possibly via lateral transport from the TTL. Analysis of

CO2 data from the Earth Resources (ER-2) high-altitude

research aircraft (Hintsa et al., 1998) at 40°N also indicated

that most of the air in the lower stratosphere must have

entered laterally from the subtropics or tropics. However,

there was also evidence that air entered the lower stratosphere directly from the midlatitude upper troposphere in

conjunction with mixing processes near the tropopause.

Several more recent observational studies support

the existence of a transition layer between the extratropical upper troposphere and lower stratosphere (see Figure

2-2). The layer exhibits chemical characteristics that are

intermediate between these two regions. For example,

ozone and carbon monoxide (CO) data collected as part of

the Stratosphere-Troposphere Experiments by Aircraft

Measurements (STREAM) program show the existence of

a transition layer that varies in thickness between winter

and summer (Hoor et al., 2001), being deeper in summer.

A study of total reactive nitrogen (NOy) and nitrous oxide

(N2O) relationships also showed that more exchange

appears to occur in the summer months. This results in a

deeper summertime transition layer with higher excess

NOy concentrations (i.e., in excess of what can be expected

from photochemical destruction of N2O) which may have

been transported into the transition layer by deep convection (Fischer et al., 2000). Another study showed that

frontal uplifting can be important for transporting trace

gases (sulfur hexafluoride (SF6), deuterated water vapor

(HDO)) from the subtropics in the troposphere into the

Arctic lower stratosphere during winter (Zahn, 2001).

These findings are consistent with seasonal variations

in exchange predicted by model studies based on wind fields

from large-scale meteorological datasets (Chen, 1995;

Haynes and Shuckburgh, 2000). In effect, exchange at

middle and high latitudes associated with synoptic eddies

and frontal systems produces a transition layer in winter and

summer that corresponds to potential temperatures between

310K and 330K. In winter at lower latitudes and higher levels

the subtropical jet acts as an effective barrier to transport

across the tropopause. Trajectory analyses point to frontal

uplifting being a more effective exchange mechanism in the

fall (Seo and Bowman, 2001) as well as the winter (Stohl,

2001). In summer the subtropical jet is a less effective barrier, and additional exchange results in a layer extending up

to ~360 K. Deep convection at midlatitudes may also be

important, but its effects are not yet well quantified.

In principle, the transition layer is a region in which

concentrations of VSL substances and their degradation

products may be relatively high and in which substantial

in situ ozone depletion may take place. Clearly, further

clarification of the mechanisms controlling the characteristics of the transition layer, the origin of air in this layer,

and its seasonal evolution are needed. Also, time scales

on which air parcels remain in this region, e.g., before

returning to the troposphere, are unknown. All these factors have implications for the lifetimes and impact of VSL

substances and their degradation products in this region.

2.2.3 Chemistry in the UT and TTL

Modeling the transport of VSL substances into the

stratosphere requires an accounting of photochemical

losses in the troposphere, particularly in the upper troposphere and tropical tropopause layer, where the VSL substances may reside for some time before ascent to the

stratosphere. This section focuses mainly on the UT.

Currently very few chemical observations of species other

than ozone are available in the TTL, which is of critical

importance for delivery of VSL substances to the stratosphere. Because the TTL is dry, sources other than water

vapor are likely to dominate the supply of odd-hydrogen

(HOx) radicals there.

Loss by photolysis in the UT can be calculated with

confidence if absorption cross section and quantum yield

data are available, because observed ultraviolet (UV)

actinic fluxes in the UT are typically within 15% of values

computed from clear-sky radiative transfer models

(Shetter et al., 2002). Loss by reaction with hydroxyl radical (OH) and possibly other radicals is more difficult to

constrain. Work over the past 5 years has profoundly

changed our understanding of HOx (OH + HO2) chemistry

in the UT (Jaeglé et al., 2001). Aircraft measurements of

HOx radicals (Wennberg et al., 1998; Brune et al., 1998,

1999; Tan et al., 2002) and supporting photochemical

models have shown that photolysis of acetone and of convected peroxides and aldehydes provides a large source

of HOx to the UT, comparable with or exceeding the

source from oxidation of water vapor (Chatfield and

Crutzen, 1984; Jaeglé et al., 1997, 1998a, 2000; Prather

and Jacob, 1997; McKeen et al., 1997; Crawford et al.,

1999; Müller and Brasseur, 1999; Collins et al., 1999;

Wang et al., 2000; Ravetta et al., 2001). Current photochemical models constrained with aircraft observations

reproduce observed HOx concentrations in the UT to

within the instrument accuracy (±40%), and account for

60-70% of the observed variance (Jaeglé et al., 2001).

Large unresolved biases remain at sunrise and sunset, in

cirrus clouds, and under high-NOx conditions (Brune et

al., 1999; Faloona et al., 2000; Jaeglé et al., 2000). The

VERY SHORT-LIVED SUBSTANCES

2.13
40% uncertainty in HOx translates into errors in the computed concentrations of some VSL species in the UT and

TTL.

Our ability to represent UT radical chemistry in

global models remains weak, despite the constraints

offered by HOx and NOx measurements from aircraft.

Limitations are not so much in our understanding of the

fast HOx photochemistry as they are in defining the factors that control the concentrations of HOx precursors and

NOx. Recognized uncertainties in models include (1) the

sources of UT NOx from lightning and from vertical transport of surface emissions, the chemical cycling between

NOx and nitric acid (HNO3), and possible complications

involving peroxynitric acid (HNO4) chemistry (Lawrence

et al., 1995; Folkins et al., 1998; Jaeglé et al., 1998b; Keim

et al., 1999; Liu et al., 1999; Folkins and Chatfield, 2000;

Schultz et al., 2000; Staudt et al., 2002); (2) the sources

and chemistry of oxygenated species (peroxides; acetone

and other carbonyls) and their deep convective transport

to the UT (Zuo and Deng, 1999; Jaeglé et al., 2000; Jacob

et al., 2002); and (3) low-temperature rate constants for

critical reactions such as methylperoxy radical (CH3O2)

plus hydroperoxyl radical (HO2) (Tyndall et al., 2001;

Ravetta et al., 2001). The possible role of halogen radicals themselves in influencing HOx chemistry, and in turn

the fate of other halogens, is another unresolved issue

(Borrmann et al., 1996; Crawford et al., 1996; Davis et

al., 1996; Solomon et al., 1997). These issues are particularly relevant in the presence of convective clouds, and

studies with cloud-resolving models (e.g., Wang and

Prinn, 2000) are providing additional insights.

2.2.4 Chemical Tracer Observations for

Model Evaluation

Observations of chemical tracers offer valuable

diagnostic information for testing atmospheric transport

of VSL substances in global models. Of particular interest

is the simulation of vertical transport from the BL to the

UT and across the tropopause. An ideal chemical tracer

for testing vertical transport in models should have the

following properties: (1) well-defined boundary conditions (concentrations or fluxes) in the BL or in the LS; (2)

well-defined production and loss rates within the atmospheric column; (3) an atmospheric lifetime such that the

transport processes of interest translate into measurable

concentration gradients; and (4) extensive observations

for model evaluation statistics. No single tracer fully

meets all these conditions, but several are sufficiently

close to be helpful.

Table 2-1 summarizes the utility of various chemical tracers for global model evaluation of the key transport processes related to simulations of VSL substances.

Current models can reproduce the vertical distributions of

radon-222 (222Rn), nonmethane hydrocarbons, and CO

without evident bias, providing some confidence in the

simulation of convective transport from the BL to the UT

at least in a global-mean sense (Jacob et al., 1997).

Simulations of ozone and 7Be indicate that models often

overestimate the downward transport of air from the LS

to the UT (Bey et al., 2001; Liu et al., 2001). Aircraft

observations of soluble gases in convective outflows suggest that the scavenging of these gases during deep convection from the BL to the UT can be successfully

described on the basis of Henry’s Law solubility and

knowledge of gas retention upon cloud freezing (Mari et

al., 2000). However, global models generally overestimate HNO3 in the UT (Thakur et al., 1999), possibly

because of insufficient scavenging, including neglect of

cirrus precipitation (Lawrence and Crutzen, 1998).

Overall, the suite of tracers in Table 2-1 provides important information for testing transport and scavenging

processes of key relevance to VSL substances, and it is

essential that any chemical transport model used to simulate VSL substances be first validated with an appropriate

ensemble of these tracers.

Ozone is a particularly valuable tracer of vertical

transport in the UT/LS region, but it has so far been underutilized for this purpose. It is the only gas sampled regularly enough in the UT (by both ozonesondes and aircraft)

to provide data representative of actual monthly means

(Logan, 1999; Thompson and Witte, 1999; Emmons et al.,

2000; Lawrence, 2001). Ozone has a lifetime against

chemical loss of months to years in the UT/LS; hence its

vertical concentration gradient is primarily determined by

transport and chemical production (Folkins et al., 1999).

Vertical profiles of ozone across the tropopause region

and correlations with water vapor provide an important

test for evaluating the nature of the tropopause and crosstropopause mass flux in global models (Law et al., 2000).

Of particular relevance for VSL substances, ozonesonde

profiles in the tropics can be used to diagnose vertical

transport in the TTL and ascent into the stratosphere

(Folkins et al., 1999). One difficulty with the use of ozone

as a tracer of vertical transport is that its long lifetime in

the UT allows for horizontal propagation of signals associated with vertical transport, and varying chemical production along these horizontal transport trajectories may

have a major effect on concentrations (Pickering et al.,

1992; Lawrence et al., 1999; Martin et al., 2002).

VERY SHORT-LIVED SUBSTANCES

2.14
2.3 ATMOSPHERIC CHEMISTRY OF

HALOGENATED VERY SHORT-LIVED (VSL)

SUBSTANCES

This section presents an assessment of the state of

knowledge of the kinetics and mechanisms of the chemical processes that determine the local lifetimes and distributions of the VSL organic halogen source gases and

their degradation products in the atmosphere. The compounds to be considered are summarized in Table 2-2.

They were selected because they have been either

observed in the atmosphere or proposed for applications

that may lead to release. Fully fluorinated VSL substances

are not covered in this Assessment, because the Fcontaining radicals produced in their degradation do

not deplete ozone.

Section 2.3.1 deals with the kinetic data and reaction mechanisms for the organic source gases, which

determine the efficiency of the SGI (source gas injection)

pathway, as defined in Section 2.1. The efficiency of the

PGI (product gas injection) pathway depends on the

kinetic data and reaction mechanisms of the degradation

products and the inorganic halogen species. These are

discussed in Section 2.3.2. Some of the species were

VERY SHORT-LIVED SUBSTANCES

2.15

Table 2-1. Chemical tracers useful for evaluating simulations of vertical transport in the troposphere,

scavenging in the troposphere, and stratosphere/troposphere exchange (STE) processes in global

models.

Tracer Application Source Sink a Selected References Pros and Cons

222Rn Vertical transport Soils Radioactive decay Jacob et al. (1997); Pros: well-defined sink, relatively

over continents (lifetime 5.5 days) Dentener et al. uniform continental source 

(1999) Cons: poor definition of source

variability, sparse observations

Nonmethane Vertical transport Anthropogenic Reaction with OH Blake et al. (1997); Pros: extensive aircraft data,

hydrocarbons, over continents and biogenic Bey et al. (2001) correlations between species

CO emissions provide additional constraints

Con: poor source definition

CH3I Vertical transport Oceans Photolysis Blake et al. (1996); Pro: extensive aircraft data

over oceans (lifetime ~5 days Bell et al. (2002) Con: poor source definition

in tropics)

210Pb Vertical transport, Decay of 222Rn Wet deposition Preiss et al. (1996); Pros: extensive surface data,

aerosol Rasch et al. good source definition

scavenging (2000) Con: sparse aircraft data

7

Be STE, subsidence Cosmic rays Radioactive decay Dibb et al. (1997); Pros: extensive surface data,

(lifetime 77 days), Liu et al. (2001) good source definition

wet deposition Con: sparse aircraft data

10Be STE (together Cosmic rays Wet deposition Dibb et al. (1994); Pro: well-defined STE tracer

with 7Be) (lifetime against Koch and Rind Con: very sparse data

radioactive decay (1998)

~106 years)

14CO STE, OH Cosmic rays Reaction with OH Jöckel et al. (1998); Pro: well-defined source distribution

(lifetime 1-3 Jöckel et al. Cons: uncertain total source

months) (2002) strength, sparse aircraft data

O3 Deep convection, Photochemistry, Photochemistry, Folkins et al. (1999); Pros: extensive sonde and aircraft

STE transport from deposition Logan (1999) data, correlation with water

stratosphere vapor provides additional

constraints

Con: complicated chemistry

Soluble gases: b Scavenging in deep Boundary layer Scavenging, Mari et al. (2000); Pros: extensive aircraft observations

CH3OOH, convection chemistry chemistry Crutzen and Cons: difficulty of separating

HCHO, H2O2, Lawrence (2000) convective transport and

HNO3 scavenging processes

a All lifetimes are given as “e-folding” lifetimes, which are the inverse of the first-order removal rate.

b CH3OOH, methyl hydroperoxide; HCHO, formaldehyde; H2O2, hydrogen peroxide; HNO3, nitric acid.
assessed in the most recent Assessment report (Kurylo

and Rodríguez et al., 1999), and the current Assessment

provides an update and extends that work to include the

inorganic halogen chemistry in the UT/LS region. The

degradation mechanism for n-propyl bromide is discussed

in Section 2.3.3. Section 2.3.4 provides the state of knowledge of iodine chemistry in the stratosphere, which is used

for the evaluation of the ozone removal efficiency of

iodine.

2.3.1 Removal of the Halogen Source

Gases

The atmospheric removal of the VSL halogen

source gases under consideration is primarily initiated via

reaction with the OH radical and/or by UV photodissociation. Removal via reactions with Cl atoms and nitrate radicals (NO3) is expected to be of only minor importance.

Table 2-3 summarizes the kinetic and photochemical

parameters for the VSL halogen source gases. For

larger molecules such as n-propyl bromide (n-PB,

CH3CH2CH2Br), the reaction rate constant with OH

depends on the site of the OH reaction (see also Section

2.3.3).

Table 2-4 summarizes estimated values of the local

lifetimes for some VSL halocarbons. For the chlorineand bromine containing compounds, reaction with OH is

the dominant loss process, apart from tribromomethane

(CHBr3) and dibromochloromethane (CHBr2Cl), where

photolysis contributes at least a comparable amount. For

alkyl iodides (such as CH3I, C2H5I, n-C3H7I, i-C3H7I,

CH2ClI, CH2BrI, and CH2I2), photolysis is always the

dominant loss process leading to prompt release of iodine

atoms:

RI + hn Æ R + I (R2.1)

Local lifetimes for these iodides can be estimated from

their absorption cross sections (Rattigan et al., 1997; Roehl

et al., 1997; Mössinger et al., 1998): 4-12 days for CH3I,

2-8 days for C2H5I, 1-3 days for i-C3H7I, 0.5-1.5 days for

n-C3H7I, hours for CH2ClI and CH2BrI, and minutes for

CH2I2. The ranges reflect variations with latitude, altitude,

and seasons. Reaction with OH may also be a significant

loss process for n-C3H7I and i-C3H7I, but I atoms are still

believed to be released promptly (Cotter et al., 2001).

VERY SHORT-LIVED SUBSTANCES

2.16

Table 2-2. VSL organic halogen source gases assessed in this report.

Source Gas Formula

Bromochloromethane a CH2BrCl

Dibromomethane (methylene bromide) a CH2Br2

Dibromochloromethane a CHBr2Cl

Bromodichloromethane a CHBrCl2

Tribromomethane (bromoform) a CHBr3

1-bromopropane (n-propyl bromide or n-PB) n-C3H7Br; CH3CH2CH2Br

Iodomethane a CH3I

Trifluoroiodomethane a CF3I 

Iodoethane (ethyl iodide) a C2H5I; CH3CH2I

1-iodopropane (n-propyl iodide) a n-C3H7I; CH3CH2CH2I

2-iodopropane (isopropyl iodide) a i-C3H7I; CH3CHICH3

Chloroiodomethane a CH2ClI

Bromoiodomethane CH2BrI

Diiodomethane a CH2I2

Dichloromethane (methylene chloride) a CH2Cl2

Trichloromethane (chloroform) a CHCl3

Trichloroethene (trichloroethylene) a C2HCl3; CHClCCl2

Tetrachloroethene (perchloroethylene) a C2Cl4; CCl2CCl2

1,2-dichloroethane C2H4Cl2; CH2ClCH2Cl

a Assessed in the 1999 Assessment (Kurylo and Rodríguez et al., 1999), which dealt only with the atmospheric observations of these species and their

interpretation. The laboratory chemical data were not assessed in detail.
VERY SHORT-LIVED SUBSTANCES

2.17

Table 2-3. Kinetic and photochemical parameters for the degradation reactions of halogenated VSL

source gases.

Species OH Rate Constant, k k300K Photolysis Rate Comments and References b

(cm3 molec–1 s

–1) (cm3 molec–1 s–1) Constant, J a

( s–1)

CH2BrCl 2.0 ¥ 10–12 exp(–870/T) 1.1 ¥ 10–13 7.7 ¥ 10–10 Sander et al. (2002)

CH2Br2 2.0 ¥ 10–12 exp(–840/T) 1.2 ¥ 10–13 <1 ¥ 10–8 Sander et al. (2002); Kurylo and

Rodríguez et al. (1999)

CHBr2Cl 5.7 ¥ 10–14 7.2 ¥ 10–8 Chiorboli et al. (1993); Kurylo and

Rodríguez et al. (1999)

CHBrCl2 1.2 ¥ 10–13 5.2 ¥ 10–8 Bilde et al. (1998); Kurylo and

Rodríguez et al. (1999)

CHBr3 1.35 ¥ 10–12 exp(–600/T) 1.8 ¥ 10–13 1 ¥ 10–6 Sander et al. (2002); Kurylo and

Rodríguez et al. (1999)

n-C3H7Br 3.1 ¥ 10–12 exp(–330/T) 1.0 ¥ 10–12 <1 ¥ 10–8 First rate constant value from

6.6 ¥ 10–18 T 2 exp(154/T) 9.9 ¥ 10–13 Sander et al. (2002), 

ka = 1.44 ¥ 10–12 exp(–450/T) ka/k = 0.32 second value and third set of

kb = 2.5 ¥ 10–18 T 2 exp(265/T) kb/k = 0.56 values from Gilles et al. 

kg = 2.9 ¥ 10–12 exp(–957/T) kg/k = 0.12 (2002) (see Figure 2-8 for

explanation of a, b, and g);

J values from Herndon et al.

(2001) and Kozlov et al. (2002)

i-C3H7Br 1.75 ¥ 10–12 exp(–260/T) 7.4 ¥ 10–13 Sander et al. (2002), Teton et al. (1996); no data

3.6 ¥ 10–12 exp(–390/T) 9.7 ¥ 10–13

CH3I 2.9 ¥ 10–12 exp(–1100/T) 7.2 ¥ 10–14 3 ¥ 10–6 Sander et al. (2002); Roehl et al. (1997)

CF3I 2.8 ¥ 10–11 exp(–2100/T) 2.4 ¥ 10–14 8 ¥ 10–6 Sander et al. (2002), Gilles et al.

(2000); Rattigan et al.

(1997), Roehl et al. (1997),

Solomon et al. (1994b)

C2H5I 3 ¥ 10–13 5 ¥ 10–6 Estimation based on correlation

7.7 ¥ 10–13 between Cl and OH

abstraction rate coefficients

from Cotter et al. (2001); Roehl et al. (1997)

n-C3H7I 1.47 ¥ 10–12 2 ¥ 10–5 Carl et al. (1998); Roehl et al. (1997)

i-C3H7I 1.22 ¥ 10–12 8 ¥ 10–6 Carl et al. (1998); Roehl et al. (1997)

CH2ClI 1 ¥ 10–4 No kinetic data; Rattigan et al. (1997),

Roehl et al. (1997)

CH2BrI 3 ¥ 10–4 No kinetic data; Mössinger et al. (1998) 

CH2I2 4 ¥ 10–3 No kinetic data; Mössinger et al. (1998),

Roehl et al. (1997)

CH2Cl2 1.9 ¥ 10–12 exp(–870/T) 1.0 ¥ 10–13 Sander et al. (2002); no data

CHCl3 2.2 ¥ 10–12 exp(–920/T) 1.0 ¥ 10–13 Sander et al. (2002); no data

C2HCl3 5.0 ¥ 10–13 exp(445/T) 2.2 ¥ 10–12 Atkinson et al. (2001); no data

C2Cl4 9.4 ¥ 10–12 exp(–1200/T) 1.7 ¥ 10–13 Atkinson et al. (2001); no data

CH2ClCH2Cl 1.05 ¥ 10–11 exp(–1142/T) 2.3 ¥ 10–13 Qiu et al. (1992); no data

a The values for the photolysis rate constants are globally and seasonally averaged estimates for 5-km altitude based on cross sections and quantum

yields measured in several laboratories (because different models were used, the values should be considered approximate).

b References cited for photolysis follow semicolon. Rate constants, cross sections, and quantum yields that have been evaluated in International Union

of Pure and Applied Chemistry (IUPAC) (Atkinson et al., 2001) and NASA Jet Propulsion Laboratory (NASA-JPL) (Sander et al., 2002) are indicated by the corresponding citations.
Deposition of the source gases into the ocean would

constitute a permanent removal if there are efficient chemical and/or biological pathways to transform the species

in the water. Estimates for the ocean removal lifetime

show that ocean removal has a negligible effect on the

local photochemical lifetime of the VSL halogen source

gases (Yvon-Lewis and Butler, 2002).

2.3.2 Production and Removal of

Degradation Products

2.3.2.1 DEGRADATION MECHANISMS FOR SOURCE

GASES

As stated, the atmospheric degradation of the

source gas is primarily initiated via reaction with OH

and/or by UV photodissociation. The radicals produced

following the initiation reaction are rapidly converted into

more stable products. Halogen atoms may be released in

the initiation reaction or during the subsequent sequence

of reactions occurring in the degradation mechanism.

The degradation mechanism for the halogencontaining compounds parallels that previously reported

for the hydrochlorofluorocarbons (HCFCs) (Cox et al.,

1995). However, many of the reactions and reaction intermediates in the degradation have not been studied or

observed in the laboratory. In those cases, formulation of

the elementary steps in the degradation mechanisms has

been based on analogy with known mechanisms. Both

OH reaction and UV photolysis loss processes will lead

to the rapid formation of alkylperoxy radicals. The degradation of the halocarbons leads to the formation of carbonyl halides, peroxynitrates, hydroxyhalomethanes, and

haloperoxy acids. The yield of a specific degradation

product will depend to some extent on the atmospheric

conditions (temperature and pressure) as well as the

atmospheric concentrations of nitric oxide (NO), nitrogen

dioxide (NO2), HO2, and organic peroxy radicals (RO2).

A general degradation mechanism for the halomethanes is shown in Figure 2-6, and a summary of the

degradation products derived from the various source

gases is given in Table 2-5. Upon photolysis, the singly

iodinated compounds (alkyl iodides) and the dihalo-iodo

compounds (CH2ClI, CH2BrI, and CH2I2) eliminate an I

atom with unit quantum yield. Photolysis therefore does

not lead to the formation of iodine-containing organic

degradation products, with the exception of CH2I2, which

could form HIC(O) as a minor product. Photolysis also

removes trifluoroiodomethane (CF3I) rapidly from the

atmosphere, with CF3 and I as photolysis products. The

atmospheric degradation of the CF3 radical does not lead

to significant ozone destruction, as discussed in Cox et al.

(1995). The photolysis of multiple-substituted halomethanes (CH2BrCl, CH2Br2, CHBr2Cl, CHBr3) will lead

to the direct elimination of a halogen atom and the formation of a formyl halide.

H-abstraction by OH is the primary loss process

for several of the less photo-labile compounds (CH2BrCl,

CH2Br2, and CH3CH2CH2Br (n-PB)) and is of secondary

importance for the iodoalkanes (C2H5I, CH3CHICH3, and

CH3CH2CH2I). For CH2BrCl and CH2Br2, upon reaction

with OH, it is likely that carbonyls such as HC(O)Br,

C(O)BrCl, and C(O)Br2 will be produced and these may

VERY SHORT-LIVED SUBSTANCES

2.18

Table 2-4. Local photochemical lifetimes (in

days) for halogenated VSL substances. The local

lifetimes are estimated using the reaction rate constant with OH and estimated photolysis rates as indicated in the table.

Compound tOH

a tJb tlocalc

CH2BrCl 150 15000 150

CH2Br2 120 5000 120

CHBr2Cl 120 d 161 69

CHBrCl2 120 e 222 78

CHBr3 100 36 26

n-C3H7Br (n-PB) 13 f >1200 f 13

i-C3H7Br 17 f g 17

CH3I 220 f 7 h (4-12) i 7

C2H5I 15-39 e, f 5 h (2-8) i 4

n-C3H7I 7.9 e, f 0.6 h (0.5-1.5) i 0.5

i-C3H7I 9.5 e, f 1.4 h (1-3) i 1.2

CH2ClI g 0.1 0.1

CH2BrI g 0.04 0.04

CH2I2

g 0.003 0.003

CH2Cl2 140 >15000 140

CHCl3 150 >15000 150

C2HCl3 4.6 f >15000 4.6

C2Cl4 100 f 1.2 ¥ 105 f 99

C2H4Cl2 70 f 70

a Calculated for T = 275 K with reaction rate constants from JPL 00-3

(Sander et al., 2000) and [OH] = 1 ¥ 106 molec cm–3 unless specified

otherwise. Values are rounded to 2 significant figures.

b Globally and seasonally averaged photolysis constant (J) value estimated for 5 km, unless specified otherwise. c From , unless specified otherwise; the

oceanic lifetime is not included (see text).

d Estimated T dependence of reaction rate (Kurylo and Rodríguez et

al., 1999). e Calculated for 300 K. 

f Calculated using OH rate constant from Table 2-3. g No data.

h Surface value.

i Range for different altitudes and latitudes.

( ) ( ) () t tt local OH

- -- = + 1 11

J
be potential carriers of bromine to the stratosphere. For

the nonmethane compounds such as CH3CH2CH2Br (nPB), reaction at the different carbon sites will lead to the

formation of different degradation products. The scheme

of the degradation mechanism for n-PB is detailed in

Section 2.3.3.

2.3.2.2 INTERMEDIATE PRODUCTS

The kinetics for the reactions involving the intermediate products are much less well understood than for

the reactions of the source gases. In most cases, this does

not matter because the initiation reaction is the ratedetermining step for the release of inorganic halogen

atoms. When there is a competition among alternative

pathways such as RO2 (where R is an alkyl or substituted

alkyl group) reactions or alkoxy radical (RO) reactions,

the lack of kinetic knowledge limits the ability to accurately predict the extent to which stable products are

formed.

The first generation of stable products following

initiation of gas-phase degradation consists generally of

carbonyl or peroxide species. Depending on the mechanism of degradation, halogen substitution may be retained

at the α or β position. Table 2-6 summarizes the data that

are known for the reactions of stable halogenated species

generated in the first steps of degradation of VSL substances. Much less is known about the site-specific

kinetics of these reactions. Generally, the presence of a

halogen atom leads to an overall reduction in the rate coefficients compared with the hydrogenated species.

Halogenated carbonyls are slightly soluble in water.

Once in solution, they undergo hydrolysis, which then

represents the driving force for their multiphase removal

from a given air mass. This chemical reaction may be

written as

RCOX (where X = Cl, Br) + H2O → RCOOH + HX

(R2.2)

The uptake coefficient, γhet, which is a convolution of all

steps involved in the multiphase processing, is expected

VERY SHORT-LIVED SUBSTANCES

2.19

CHXYZ X,Y = H, Cl, Br, I

Z = Cl, Br, I

CXYW

OH/Cl

O2

OOCXYW

NO

RO2

OCXYW

C(O)XY + W X,Y,W ≠ H

C(O)HX + Y

H2C(O) + W X,Y = H

W = H

∆

NO2

O2NOOCXYW

∆, hν

hν

HO2

HOOCXYW

OH

hν

OH 

HOOC(O)H X ,Y = H

HOOC(O)X HOOC(O)Y

hν

hν

W = Cl, Br, I for reaction

W = H for hν

RO2 /HO2 C(O)YW (X = H atom)

Multiphase

Processing

Multiple Steps

Minor channels

X = H

Figure 2-6. Halomethane degradation mechanism. Stable intermediate products containing halogens are

enclosed in solid boxes. Minor pathways and products are indicated by dashed arrows and boxes.
to be small for all compounds. Current estimates give an

upper limit of 10–3 (de Bruyn et al., 1995; George et al.,

1994), which indicates that cloud removal may play a role

if gas-phase reactions are slow. It must also be emphasized that, for the compounds with very low reactivity,

cloud evaporation may re-inject these compounds into the

air, possibly at a different altitude, before multiphase processing can occur.

2.3.2.3 INORGANIC BROMINE AND IODINE SPECIES

Br and I atoms released in the troposphere through

the degradation of the source gases and their degradation

products undergo a series of reactions in which they are

interconverted among the oxides (XO) and inorganic

reservoirs: HX (hydrogen halides), XONO2 (halogen

nitrates), and HOX (hypohalous acids), where here X =

Br or I. The partitioning reactions for the inorganic reservoirs are common to bromine and iodine, and are illustrated in Figure 2-7. Bromine does not form higher oxides

but iodine monoxide (IO) can react with bromine

monoxide (BrO) or IO at a rate competitive with HO2 to

form iodine dioxide (OIO) (Cox et al., 1999; Rowley et

al., 2001). The chemical reactions of OIO are not well

known. Photodissociation to the ground state of atomic

oxygen (O(3

P)) + IO is not favored thermodynamically in

its main visible absorption band (Misra and Marshall,

1998) and is slow (Ingham et al., 2000). This photolysis

VERY SHORT-LIVED SUBSTANCES

2.20

Table 2-5. Summary of the degradation products from halogenated VSL substances.

Source Gas Major Degradation Products Minor Degradation Products

CH2BrCl HClC(O), inorganic HClBrC(OOH)

bromine

CH2Br2 HBrC(O), inorganic HBr2C(OOH)

bromine

CHBr2Cl HClC(O), ClBrC(O), HBr2C(OOH),

inorganic bromine ClBr2C(OOH)

CHBrCl2 HClC(O), Cl2C(O), Cl2BrC(OOH),

inorganic bromine HCl2C(OOH)

CHBr3 HBrC(O), Br2C(O) Br3C(OOH),

HBr2C(OOH)

n-C3H7Br (n-PB) CH3C(O)CH2Br, Hydroperoxy compounds,

inorganic bromine, peroxycarbonyl nitrates,

CH2BrCH2C(O)H peroxynitrates

CH3I Inorganic iodine —

CF3I Inorganic iodine —

C2H5I Inorganic iodine CH2IC(O)H

n-C3H7I Inorganic iodine CH3C(O)CH2I,

CH2ICH2C(O)H

i-C3H7I Inorganic iodine CH3CHIC(O)H

CH2ClI Inorganic chlorine and iodine H2ClC(OOH)

CH2BrI Inorganic bromine and iodine H2BrC(OOH)

CH2I2 Inorganic iodine HIC(O),

H2IC(OOH)

CH2Cl2 HClC(O) HCl2C(OOH)

CHCl3 Cl2C(O) HCl2C(OOH)

C2HCl3 Cl2C(O), HClC(O) HClC(OH)C(O)Cl,

HClC(OH)C(OOH)Cl2,

Cl2C(OH)C(O)H,

HClC(OOH)C(OH)Cl2

C2Cl4 Cl2C(O), ClC(O)OH Cl2C(OH)CCl2(OOH)

CH2ClCH2Cl HClC(O), inorganic CH2ClCHCl(OOH)

chlorine
VERY SHORT-LIVED SUBSTANCES

2.21

Table 2-6. Kinetic and photochemical parameters for the gas-phase reactions of the degradation products of halogenated VSL substances.

Species Radical a k k300 Photolysis Rate Comments and References c

(cm3 molec–1 s

–1) (cm3 molec–1 s–1) Constant, J b

(s-1)

Cl2C(O) OH <5.0 ¥ 10–15 2 ¥ 10–9 Atkinson et al. (1997);

Cox et al. (1995)

Br2C(O) No kinetic or

photochemical data

BrClC(O) No kinetic or

photochemical data

HClC(O) OH <5.0 ¥ 10–13 1 ¥ 10–8 Atkinson et al. (1997)

(limited experimental

data); Cox et al. (1995)

HClC(O) Cl 1.2 ¥ 10–11 exp(–820/T) 7.9 ¥ 10–13 Atkinson et al. (1997)

HBrC(O) 3 ¥ 10–6 No kinetic data; estimate 

CH3C(O)CH2Br 1 ¥ 10–4 No kinetic data;

Burkholder et al. (2002)

CH2ClC(O)H OH 3.1 ¥ 10–12 Atkinson et al. (1997)

CCl3C(O)Cl No kinetic or

photochemical data

HOI 1.2 ¥ 10–2 No kinetic data;

Bauer et al. (1998)

HOBr OH <5 ¥ 10–13 3 ¥ 10–3 Kukui et al. (1996), upper

limit; Ingham et al. (1998)

HOBr Cl 8 ¥ 10–11 Kukui et al. (1996), products

are BrCl and OH

CH2ClOOH Cl d 5.9 ¥ 10–13 Wallington et al. (1996),

no kinetic data for OH

reactions and photolysis

CCl2BrOOH No kinetic or

photochemical data

BrONO2 ~1 ¥ 10–3 Burkholder et al. (1995)

IONO2 4 ¥ 10–2 Mössinger et al. (2002) 

OIO <10–2 No kinetic data; Ingham et al. 

(O(3P) channel) (2000) (upper limit for O(3P)

0.3-2.2 channel), Ashworth et al.

(I + O2 channel) (2002) (products are I + O2)

a Cl atom reactions included if kCl > 50 ¥ kOH or no OH kinetic data available.

b Atmospheric photolysis rate constants are globally and seasonally averaged estimates for 5-km altitude (because different models were used, the

values should be considered approximate). c References cited for photolysis follow semicolon.

d Strongly deactivating compared with reaction of Cl + CH3OOH (k300 = 5.9 ¥ 10–11 cm3 molec–1 s

–1) (Atkinson et al., 1997).
VERY SHORT-LIVED SUBSTANCES

2.22

pathway would lead to a null cycle for ozone. However,

photolysis to I + O2 may occur (Ashworth et al., 2002),

leading to an ozone loss cycle:

IO + BrO → Br + OIO 

OIO + hν → I + O2

I + O3 → IO + O2

Br + O3 → BrO + O2

2 O3 → 3 O2

Thus OIO is a potentially significant atmospheric iodine

species, and indeed has been detected in the marine

boundary layer (Allan et al., 2001). Further oxidation of

OIO, either by gas-phase or multiphase processes, to IO3

–

is a likely sink for tropospheric iodine, leading to the iodate

observed in precipitation (Truesdale and Jones, 1996).

The relative importance of the halogen nitrates and

HOX reservoirs depends on the local NO2 and HO2 concentrations, and on the presence of an aqueous phase where

halogen nitrates may be hydrolyzed to form HOX. The

inorganic reservoirs are removed from the troposphere by

incorporation into aerosols, cloud droplets, and ice particles (cirrus clouds in the UT). However, the incorporation

of the reservoirs into aerosols may not lead to permanent

loss. Laboratory studies (Abbatt and Waschewsky, 1994;

Fickert et al., 1999; Holmes et al., 2001; Mössinger and

Cox, 2001) show that HOX and XONO2 species react very

efficiently with Cl– and Br– on aerosols and ice particles,

releasing the corresponding halogen or inter-halogen molecules, which all photodissociate to atomic halogens:

HOBr (or HOI) + Cl– + H+ → BrCl (or ICl) + H2O

(R2.3)

HOBr (or HOI) + Br– + H+ → Br2 (or IBr) + H2O

(R2.4)

Reactions (R2.3) and (R2.4) are very efficient in acidic

solutions or on HCl or hydrogen bromide (HBr)-doped

surfaces. These reactions (and similar reactions involving

XONO2) have been the focus of many studies, especially

in connection with their role in stratospheric halogen activation. They also play a role in tropospheric chemistry,

particularly in the marine environment where sea salt

aerosols contain Cl– and Br– ions. Additional reactive

halogen is produced autocatalycally, leading for example

to the so-called “bromine explosion” observed in the Arctic

(e.g., Barrie and Platt, 1997). To be efficient, these reactions require cold conditions (such as frozen aerosols in

the Arctic troposphere) and sustained levels of particles

(such as ice or sulfuric acid aerosols in the stratosphere)

that may be observed, for instance, under volcanic influence. Table 2-7 lists the required parameters for an assessment of the fate of some of the HOX and XONO2 species.

Reactions similar to (R2.3) and (R2.4) efficiently

recycle iodine and bromine into radical forms, thereby

effectively increasing the lifetime of tropospheric halogens. This explains the relatively high abundance of IO

in the marine boundary layer. There are also recent direct

observations that show that BrO may be present in the

upper troposphere at a concentration of 0.4-2 parts per

trillion (ppt), which is much higher than predicted on the

basis of known photochemical sources and heterogeneous

removal rates (Fitzenberger et al., 2000). Such observations indicate that the recycling processes may be important, leading to a significant increase of the efficiency of

the product gas injection (PGI) pathway in supplying

inorganic halogen species to the stratosphere.

2.3.3 Degradation Mechanism for n-Propyl

Bromide (n-PB)

The compound 1-bromopropane (or n-propyl bromide (n-PB), CH3CH2CH2Br) has been suggested for use

in a variety of applications that would lead to its release

into the troposphere. A reliable estimate of the amount of

inorganic bromine that will be delivered to the strato-

? Multistep Oxidation

hν

(IO3

-

)adsorbed

HO

NO2

HO2

RH, HO2

 ∆

hν

 ?

∆

hν

hν

(H++X-

)adsorbed

OIO

I2O2 XO XONO2

HX X HOX

NO

O3

X = Br or I

IO

XY

Yheterogeneous

 IO

not X=Br

Y= Cl or Br

hν

Figure 2-7. Inorganic halogen tropospheric chemistry.
VERY SHORT-LIVED SUBSTANCES

2.23

sphere from release of n-PB requires better knowledge of

its degradation mechanism and a guess as to the possible

future quantities produced and emitted.

The degradation mechanism for n-PB is outlined

in Figure 2-8. Recent laboratory measurements have

determined the site-specific OH rate coefficients over the

temperature range 230 to 360 K (Gilles et al., 2002). The

kinetic data are reported in Table 2-3. The reaction rate

constants reported by Gilles et al. (2002) and that recommended in Sander et al. (2002) give estimated local lifetimes of 13 days (Table 2-4).

Bromoacetone (CH3C(O)CH2Br) was observed as

a major degradation product following the OH reaction at

the -CH2- site of the n-PB molecule (b channel in Figure

2-8). Burkholder et al. (2002) have recently shown that

CH3C(O)CH2Br will rapidly photolyze in the atmosphere

(~hours). The photolysis products were not determined,

but the possible bromine-containing products are expected

to liberate bromine within days. The reaction of OH at

the -CH2Br group in CH3CH2CH2Br (a channel) will lead

to the rapid elimination of Br and the formation of

CH3CH2C(O)H. The reaction of OH at the CH3- group in

CH3CH2CH2Br (g channel) will predominantly lead to the

formation of bromoaldehydes, (O)CHCH2CH2Br and

(O)CHCH2Br. These compounds are expected to degrade

in the atmosphere (~days) through reaction with OH and

UV photolysis. Minor reactions of the reaction intermediates can lead to the formation of peroxynitrates (containing the OONO2 group), peroxycarbonylnitrates

(PAN-type compounds, resulting from the further oxidation of the bromoaldehydes), and alkyl hydroperoxides

(containing the OOH group). The peroxynitrates will

thermally decompose rapidly under all tropospheric

conditions. The alkyl hydroperoxides will be removed

via a combination of OH reaction, UV photolysis, and

multiphase processing. The rates for these processes are

not known but are expected to lead to atmospheric lifetimes on the order of days. The only possible degradation

products of n-PB with atmospheric lifetimes longer than

the parent compound are the peroxycarbonylnitrates, but

Table 2-7. Data for the multiphase chemistry of some inorganic degradation products. The column for

ghet gives the reaction probability per collision with the surfaces.

Molecule Henry’s Reaction ghet Temperature References

Law Constant (K)

(M atm–1)

HOBr >1900 Water >0.01 274 Blatchley et al. (1992)

(hypo- NaCl 6.5 ¥ 10–3 300 Atkinson et al. (2001)

bromous acid) NaBr 0.18 300 Atkinson et al. (2001)

HCl-doped ice 4 ¥ 10–3 180-222 Chu and Chu (1999)

to 0.3

Sulfuric acid >0.01 238 Waschewsky and Abbatt (1999)

BrONO2 >7.1 ¥ 104 Water 0.06 >273 Aguzzi and Rossi (1999)

(bromine NaCl 0.3 – Aguzzi and Rossi (1999)

nitrate) NaBr 0.3 – Aguzzi and Rossi (1999)

HCl-doped ice >0.2 200 Atkinson et al. (2001)

Sulfuric acid >0.2 294 Atkinson et al. (2001)

HOI >45 Water – – Thompson and Zafiriou (1983)

(hypoiodous NaCl 1.6 ¥ 10–2 278-298 Atkinson et al. (2001)

acid) NaBr 3.4 ¥ 10–2 278-298 Mössinger and Cox (2001)

HCl-doped ice 5 ¥ 10–2 200 Holmes et al. (2001)

Sulfuric acid 1.5 ¥ 10–2 195-252 Holmes et al. (2001)

to 7 ¥ 10–2

IONO2 – Water – – –

(iodine NaCl >0.1 253 Holmes et al. (2001)

nitrate) NaBr >0.1 253 Holmes et al. (2001)

HCl-doped ice >0.05 243 Holmes et al. (2001)

Sulfuric acid >0.3 253 Holmes et al. (2001)
VERY SHORT-LIVED SUBSTANCES

2.24

only if they are formed, if they contain Br, and if they are

rapidly transported into the low-temperature environment

of the upper troposphere. In conclusion, it is very likely

that the product gas injection pathway for n-PB is from

transport of inorganic halogen rather than intermediate

products.

2.3.4 Stratospheric Iodine Chemistry

Although the chemistry of chlorine and bromine in

the stratosphere is reasonably well known, this is not the

case for iodine. Chameides and Davis (1980) were the

first to describe the atmospheric photochemistry of iodine

in the marine boundary layer, where the iodine is produced

from the oceanic source of volatile alkyl iodides. If alkyl

iodides or their degradation products could enter the stratosphere, the I atom would be rapidly converted to IO

through reaction with O3. Recycling of IO to I through

reactions with ClO, BrO, and HO2 would constitute catalytic O3 removal cycles.

Solomon et al. (1994a) provided the first estimate

for the ozone removal efficiency of iodine in the stratosphere. In their study, inorganic iodine was partitioned

among I, IO, hydrogen iodide (HI), hypoiodous acid

(HOI), and iodine nitrate (IONO2), with 90% of the iodine

in the form of I and IO. The ozone removal cycles were

identified as

IO + HO2 → HOI + O2

HOI + hν → OH + I

I + O3 → IO + O2

OH + O3 → ΗΟ2 + Ο2

2 Ο3 → 3 O2

and

CH3CH2CH2Br

OH

O2

NO 

RO2

∆

OCH2CH2CH2Br

CH3CH(O)CH2Br

CH3CH2C(O)HBr

CH3CH2C(O)H + Br

(O)CHCH2CH2Br

O2

O2

CH3C(O)CH2Br

∆ H2C(O) + CH2CH2Br

Multiple Steps

(O)CHCH2Br

NO2

∆, hν

hν

HO2

OH 

hν hν OH 

Br-Carbonyl Compounds 

and Br elimination

Multiple Steps

Minor channels

Products + Br

RO2/HO2

Multiphase 

Processing

CH2CH2CH2Br γ

CH3CHCH2Br β

CH3CH2CHBr α

Branching Ratio

CH3CH=CH2 + Br

(OO)CH2CH2CH2Br

CH3CH(OO)CH2Br

CH3CH2C(OO)HBr

CH3CH(OONO2)CH2Br

CH3CH2C(OONO2)HBr

Products

(OOH)CH2CH2CH2Br

CH3CH(OOH)CH2Br

CH3CH2C(OOH)HBr

∆

(OONO2)CH2CH2CH2Br

Figure 2-8. Degradation mechanism of n-propyl bromide (CH3CH2CH2Br). The major channel (β, in green)

leads to the formation of bromoacetone (CH3C(O)CH2Br) (see text for details). Dashed and solid boxes as for

Figure 2-6.
VERY SHORT-LIVED SUBSTANCES

2.25

IO + XO Æ I + X + O2

(or IX + O2, followed by IX + hn Æ I + X)

I + O3 Æ IO + O2

X + O3 Æ XO + O2

2 O3 Æ 3 O2

where here X is Br or Cl. At the time the study was done,

rate data were not available for the reactions of IO with

BrO and ClO. Solomon et al. (1994a) assumed a rate of

1 ¥ 10–10 cm3 molec–1 s–1 for both rate coefficients and

estimated that the ozone removal efficiency for iodine is

about 1000 times that of chlorine. It was suggested that

the coupling of a low natural abundance of iodine species

with increasing bromine and chlorine levels has contributed to the trend in ozone loss in the lowermost stratosphere.

Laboratory work in the past 5 years has significantly improved the knowledge of the relevant kinetic

data. For the IO + ClO reaction, the rate coefficient has

been found to be much lower than estimated by Solomon

et al. (1994a), by a factor between 8 at 298 K and 5 at temperatures of the lower stratosphere (Bedjanian et al., 1997;

Turnipseed et al., 1997). Furthermore, three product channels have been identified:

(a) IO + ClO Æ I + OClO

(b) Æ I + Cl + O2

(c) Æ ICl + O2

Channel (a), which leads to a null cycle for ozone depletion, is the major channel: ka/(ka+kb+kc) = 0.55. The channels that lead to ozone loss, (b) and (c), account for a significant fraction of the reaction: (kb+kc)/(ka+kb+kc) = 0.45

at 298 K (Bedjanian et al., 1997; Turnipseed et al., 1997).

Several studies reported rate coefficients for the

IO + BrO reaction ranging from 7 ¥ 10 –11 to 10 ¥ 10 –11

cm3 molec–1 s–1 (Gilles et al., 1997; Laszlo et al., 1997;

Bedjanian et al., 1998; Rowley et al., 2001). The possible

channels for this reaction are:

(a) IO + BrO Æ Br + OIO

(b) Æ I + Br + O2

(c) Æ IBr + O2

(d) Æ I + OBrO

The branching ratio for channel (a) here has been measured to be >0.65 (Bedjanian et al., 1998) while channels

(b) and (c) account for less than 0.35 of the reaction at 298

K (Gilles et al., 1997; Bedjanian et al., 1998). The OIO

molecule formed in channel (a) has been unambiguously

detected in this reaction (Cox et al., 1999). As discussed

in Section 2.3.2.3, photolysis of OIO would efficiently

produce I + O2, and channel (a) could lead to an ozone

loss cycle. However, uncertainties in OIO chemistry do

not allow any firm conclusions regarding the ozone depletion efficiency of iodine in the stratosphere. The new

data obtained for the IO + IO reaction are not discussed

here, because this reaction, which also produces OIO, is

not relevant to stratospheric iodine chemistry, but could

be important in the marine boundary layer.

Laboratory investigations of the temperature

dependence of the IO + HO2 reaction (Maguin et al., 1992;

Canosa-Mas et al., 1999; Cronkhite et al., 1999; Knight

and Crowley, 2001) lead to a somewhat higher rate coefficient (up to a factor 3 at stratospheric temperatures) than

that used in the calculations of Solomon et al. (1994a).

This, associated with the reduced effect of the reactions

of IO with both ClO and BrO, will increase the relative

importance of the HO2 cycle in the ozone loss due to

iodine. Finally, in all reactions of IO with other radicals,

the existence of a termolecular channel (similar to the ClO

dimer-forming channel for ClO + ClO) was not observed.

The laboratory data in Sander et al. (2000) have

been used for a recent re-evaluation of the ozone removal

efficiency of iodine relative to chlorine; a value of 150-

300 is used in this chapter (see Section 2.5) based on estimates provided by Bedjanian et al. (1998) and simulations

using the Atmospheric and Environmental Research, Inc.,

model (see Ko et al., 1998). However, these estimates,

which are lower than the value of 1000 initially suggested

by Solomon et al (1994a), are still very uncertain because

of the lack of knowledge about OIO chemistry, and also

about the heterogeneous reactions that could be very

important for iodine, since all the inorganic reservoirs are

reactive toward stratospheric particles (both ice and sulfuric acid (H2SO4) surfaces). Many data are still missing

for these heterogeneous processes, such as those needed

to estimate the hydrolysis rate of IONO2. When such data

are available, a multidimensional atmospheric model

should be used to perform a more rigorous calculation of

the ozone removal efficiency of iodine.

2.4 CONTRIBUTION OF HALOGENATED VSL

SUBSTANCES TO THE STRATOSPHERIC

INORGANIC HALOGEN BUDGET

There are two independent ways to obtain the total

local concentration of inorganic halogen species in the

present-day stratosphere. The first method (inorganic

method) is to add up the measured concentrations of each

inorganic species (HCl, ClONO2, HOCl, ClO, Cl2O2, etc.,

in the case of chlorine). A variant of this method uses
VERY SHORT-LIVED SUBSTANCES

2.26

measured concentrations of a subset of the inorganic

halogen species (e.g., BrO) and derives the other concentrations using a photochemical model. The second method

(the organic method) is to add the contributions from each

source gas by taking the difference between the measured

concentration of the source gas and its expected concentration in an air parcel based on when it first entered the

stratosphere (see, e.g., Prather and Watson, 1990; Woodbridge et al., 1995; Wamsley et al., 1998). Contributions

from transport of very short-lived (VSL) source gases,

degradation products, and inorganic halogen species from

the troposphere are usually ignored in the organic method.

Agreement between the two methods provides some reassurance, but not irrefutable, that our understanding of the

present-day halogen budget may be correct. For example,

Woodbridge et al. (1995) carried out a study of the inorganic chlorine budget and concluded that inorganic chlorine (Cly) concentrations derived using the two methods

are in reasonable agreement. However, the work of

Wamsley et al. (1998) showed that derived inorganic bromine (Bry) using the organic method is 30% smaller than

the expected Bry from the measured BrO, but that these

differences are barely within calculated uncertainties.

Recent analyses by Pfeilsticker et al. (2000) of a different

dataset concluded that the two methods give values that

agree within the estimated uncertainties (approximately

±3 ppt) (see Chapters 3 and 4).

The measured concentrations of VSL source gases

such as CH3I and CHBr3 are sufficiently high near the

tropical tropopause that they could make a non-negligible

contribution (see Table 2-8; review by Kurylo and

Rodríguez et al., 1999; Sturges et al., 2000). This section

summarizes the observations of the VSL halogen source

gases and provides an estimate of their contribution to the

present-day halogen budget in the stratosphere. Recent

modeling studies are also reviewed and provide some estimates of the amount of Bry entering the stratosphere from

particular substances (CHBr3, n-PB) and their degradation products.

2.4.1 Observed Concentrations and

Sources of Organic Chlorine,

Bromine, and Iodine Compounds

The atmospheric distribution and sources of

organic chlorine, bromine, and iodine compounds were

reviewed in Kurylo and Rodríguez et al. (1999). Recent

measured concentrations of the species in the tropical troposphere from Transport and Chemical Evolution over

the Pacific (TRACE-P) and Pacific Exploratory Mission

in the Tropical Pacific (PEM-Tropics) A and B are summarized in Table 2-8 and compared with values given in

the previous Assessment (Kurylo and Rodríguez et al.,

1999). Burdens estimated using the numerator in the

right-hand side of Equation (2.3) in Section 2.4.2 are given

in Table 2-9. The estimated local lifetimes from Table 2-

4 were used to calculate the sources for each species.

These values were compared with previous estimates

found in the literature. More recently, measurements

taken from bottle samples collected on commercial aircraft as part of the European Civil Aircraft for Regular

Investigation of the Atmosphere Based on an Instrument

Container (CARIBIC) project (http://www.caribic-atmospheric.com/) have also been analyzed for concentrations

of halogen-containing source gases. These will prove useful in future studies. Reported measurements of degradation products or inorganic species in the troposphere are

scarce, but include OIO (Allan et al., 2001), BrO

(Fitzenberger et al., 2000), and phosgene (Kindler et al.,

1995).

2.4.1.1 CHLORINE COMPOUNDS

Current measurements indicate that the mixing

ratio of inorganic chlorine is about 3 parts per billion (ppb)

in the upper stratosphere (based on measurement of HCl)

and about 1.5 ppb in the lower stratosphere (based on the

organic method). The VSL chlorine compounds have freetropospheric concentrations in the range 1-50 parts per

trillion (ppt). Many of them have an anthropogenic component in their sources. Methylene chloride (dichloromethane, CH2Cl2), trichloroethylene (trichloroethene,

C2HCl3), and perchloroethylene (tetrachloroethene,

C2Cl4) are mainly used as industrial solvents, and 1,2-

dichloroethane (C2H4Cl2) is used as a lead scavenger in

leaded gasoline. Concentrations of these gases in some

urban areas reach ppb levels but their concentrations in

the free troposphere are in the 1-30 ppt range (Blake et

al., 1999a, 2001). Khalil et al. (1999) and references

therein also suggested that natural emissions of chloroform (CHCl3), predominantly from oceans, are important

sources. This VSL also appears to have other small

sources such as pulp and paper processing and water treatment. Biomass burning is thought to be a very minor

source for CHCl3. Many of the VSL chlorine source gases

exhibit significant seasonal and latitudinal variations in

their concentrations (Wang et al., 1995; Yokouchi et al.,

1996; Blake et al., 1999a). This could be due to variations in the emissions and variations in OH, because reaction with OH is the major removal process. Phosgene

(COCl2) is not emitted directly to the atmosphere but is a

breakdown product of methyl chloroform and some of the

gases listed above (Kindler et al., 1995). Its atmospheric

concentration is in the 10-60 ppt range and its lifetime
VERY SHORT-LIVED SUBSTANCES

2.27

Table 2-8. Estimated values for (median tropical surface (<1 km) mixing ratios) and (median

mixing ratios in the tropical upper troposphere (8-12 km)). The data for TRACE-P were collected over the

Western Pacific and the South China Sea, 23.5°N to 8°N. The data for PEM-Tropics A and B were from over

the remote Pacific, 15°N to 15°S.

C2HCl3 NH, 1-5; 0-0.1 1 (0.3-2.0) 0.2 (0.02-0.28) 0.2

SH, 0.01-0.1 N/A N/A N/A

N/A N/A N/A

CH3I 0.1-2.0 0.05-0.2 0.6 (0.5-0.8) 0.03 (0.02-0.08) 0.05

0.42 (0.3-0.5) 0.09 (0.06-0.12) 0.2

0.445 (0.4-0.6) 0.061 (0.04-0.09) 0.15

CHBr3 0.6-3.0 0.4-0.6 1.1 (0.9-1.5)) 0.3 (0.2-0.5) 0.3

1.82 (1.0-2.3) 0.54 (0.4-0.7) 0.3

1.83 (1.4-2.2) 0.42 (0.3-0.5) 0.2

CH2ClCH2Cl NH, 20-40; 14.9 ± 1.1 8.8 (5.7-14.5) 2.9 (2.2-3.3) 0.3

SH, 5-7 1.53 (0.8-1.7) 0.92 (0.7-1.3) 0.6

N/A N/A N/A

CHBr2Cl 0.1-0.5 0.04-0.11 0.07 (0.06-0.08) 0.04 (0.03-0.05) 0.6

0.20 (0.14-0.26) 0.10 (0.08-0.12) 0.5

0.175 (0.14-0.21) 0.08 (0.06-0.11) 0.5

CHBrCl2 0.12-0.6 0.04-0.11 0.19 (0.17-0.22) 0.14 (0.12-0.15) 0.7

0.21(0.14-0.26) 0.13 (0.11-0.14) 0.6

0.17 (0.15-0.20) 0.08 (0.05-0.11) 0.5

C2Cl4 NH, 5-15; 1-3 6.5 (4.5-10.3) 2.1 (1.7-2.5) 0.3

SH, 0.7-1.5 2.34 (1.5-2.9) 1.38 (1.2-1.7) 0.6

1.3 (1.2-1.16) 1.4 (1.2-1.6) 1.1

CH2Br2 0.8-3.4 0.6-0.9 1.15 (1.0-1.2) 0.9 (0.7-1.0) 0.8

1.33 (1.1-1.5) 0.92 (0.85-1.0) 0.7

1.07 (0.93-1.18) 0.81 (0.77-0.85) 0.8

CH2Cl2 34 (27-39) 17 (16-19) 0.5

13.9 (9-16) 10.1 (9-12) 0.7

11.4 (10.9-13) e 13.1 (12.4-14) e 1.1

CHCl3 NH, 10-15; 3.1 ± 0.7 11.4 (9.0-13.3) 6.9 (6.3-7.5) 0.6

SH, 5-7 6.6 (5.2-7.1) 5.1 (4.8-5.6) 0.8

6.0 (5.7-6.4) 5.8 (5.6-6.1) 1.0

In columns 4-6, for each species, the first value is for data from TRACE-P; the second, from PEM-Tropics B; and the third, from PEM-Tropics A. N/A

denotes data not available.

a From Kurylo and Rodríguez et al. (1999).

b Range is 25th-75th quartile. c The results for PEM-Tropics A and B can be found in Blake et al. (1999a,b, 2001, 2002). For a description of the campaigns and data availability, go

to http://www-gte.larc.nasa.gov/gte_fld.htm. 

d Number of samples. e Data from a subset of cans (BL, n = 89; UT, n = 41) analyzed by E. Atlas, NCAR, U.S.

f X

UT f X

BL

Species Estimated Estimated

Tropospheric 10-km

Mixing Tropical

Ratio a Mixing

(pptv) Ratio a

(pptv)

and Range b and Range from: b from: Estimated from:

TRACE-P c (331) d TRACE-P c (280) d

PEM-Tropics B c (572) d PEM-Tropics B c (433) d

PEM-Tropics A c (337) d PEM-Tropics A c (241) d

(pptv) (pptv)

f X

UT f X

BL

f

f

X

X

UT

BL

TRACE-P

PEM-Tropics B

PEM-Tropics A
VERY SHORT-LIVED SUBSTANCES

2.28

Table 2-9. Estimated local lifetimes, burdens, removal rates, and sources for some halogenated VSL

compounds.

Compound tlocal Estimated Estimated Estimated Source from

(days) Burden Removal Rate Inventory or Estimate of

(Gg) (Gg yr–1) Biogeochemical Cycle 

(Gg yr–1)

CH2BrCl 150 a 1.2 (Br) *1 2.9 (Br) c

CH2Br2 120 a 18-22 (Br) b 55-67 (Br) c

120 a 19 (Br) *1 58 (Br) c

CHBrCl2 78 a 1.3-1.5 (Br) b 6.1-7.0 (Br) c

78 a 1.2 (Br) *1 5.5 (Br) c

CHBr2Cl 69 a 0.8-2.2 (Br) b 4.2-12 (Br) c

69 a 2.3 (Br) *1 12 (Br) c

CHBr3 26 a 11-18 (Br) b 150-250 (Br) c

26 a 14 (Br) *1 200 (Br) c 209 (47-370) (Br) *3

285 *2, d

CH3I 5 e 1.7-2.2 b (I) 120-160 (I) c 90-450 (I) *4

6 f 4.8 (I) f 214 (I) f

C2H5I 4 a 0.5 (I) 46 (I) c

CH2Cl2 140 a 83-250 (Cl) b 220-650 (Cl) c 583 (or 487 (Cl))*5 (industrial)

140 a 250 (Cl) *6 650 (Cl) c 160 (Cl) ocean *6

180 *6, e 250 (Cl) *6 500 (Cl) *6, d 49 (Cl) biomass burning *6

CHCl3 150 a 66-130 (Cl) b 160-320 c 564 (Cl) *6

150 a 210 (Cl) *7 511 c

183 *7, e 210 (Cl) *7 470 (350-600) *7, e

412(Cl) *6, d

C2Cl4 99 a 17-85 (Cl) b 63-310 (Cl) c 366 (or 313 (Cl))*5 (industrial)

99 a 160 (Cl) *6 590 (Cl) c 16 (Cl) ocean *6

133 *6, e 160 (Cl) *6 440 (Cl) *6, e 2 (Cl) fossil fuel *6

C2HCl3 4.6 a 3.1 (Cl) b 260 (Cl) c 241 (or 195 (Cl)) *5 (industrial)

4.6 a 5.30 (Cl) *6 440 (Cl) c 20 (Cl) ocean *6

5.5 *6, e 5.30 (Cl) *6 350 (Cl) *6, e 3 (Cl) fossil fuel *6

COCl2 70 g 250 (Cl) *1 1000 c

CH2ClCH2Cl 70 a 5-26 (Cl) b 26-130 (Cl) c 700 *8

References: *1 Kurylo and Rodríguez et al. (1999); *2 Dvortsov et al. (1999); *3 Carpenter and Liss (2000); *4 Singh et al. (1983), Liss and Slater

(1974), Moore and Groszko (1999); *5 McCulloch et al. (1999); *6 Keene et al. (1999); *7 Khalil et al. (1999); *8 Khalil (1998).

a From Table 2-4.

b Burden estimated using Equation (2.3) with the median BL mixing ratios from Table 2-8 and estimated scale height given in Table 2-10. The range

reflects the three different median mixing ratios from TRACE-P, PEM-Tropics A, and PEM-Tropics B. c From estimated lifetime and estimated burden, i.e., (column 3)/(column 2). The range reflects range in estimated burden.

d From 2-D model. e From 12-box model as calculated in reference 3.

f From 3-D model simulation from Bell et al. (2001). g Wet-removal lifetime.
from wet deposition is about 2 months. Textor et al. (2002)

revisited the issue of volcanic injection of chlorine to the

stratosphere. However, reliable numerical estimates are

not yet available.

2.4.1.2 BROMINE COMPOUNDS

Based on measured concentrations of BrO in the

lower stratosphere, the concentration of Bry is estimated

to be about 20 ppt. The observed concentrations of VSL

bromocarbons in the BL are in the 0.02-5 ppt range. Most

have higher concentrations in the marine BL in regions

with high ocean productivity. This indicates a possible

marine source, although the exact mechanism for production is not well understood in most cases. Dibromomethane (CH2Br2; methylene bromide) is reasonably well

mixed throughout the troposphere, compared with other

bromocarbons, in part because of its widespread marine

emissions and also because its lifetime is about 4 months

(Table 2-4). Blake et al. (2002) reported mixing ratios of

0.8-1.2 ppt to altitudes of 12 km in tropical regions, and

Schauffler et al. (1999) observed concentrations of 0.3-

0.7 ppt at altitudes of 15-19 km in the tropics. The lower

values from Schauffler et al. (1999) are consistent with

the fact that Stratospheric Tracers of Atmospheric

Transport (STRAT) data were obtained in the regions and

seasons when convection was less prevalent. Bromoform

(CHBr3) is mainly emitted from marine source(s), but it

has recently been reported that it may also have a minor

source from industrial processes (Blake et al., 2002).

Photolysis dominates its removal (Table 2-4). Blake et

al. (2002) reported average marine boundary layer (MBL)

mixing ratios of 0.7-1.9 ppt from their Pacific airborne

samples collected between 35°N and 50°S. Higher marine

boundary layer concentrations were observed in equatorial regions. Average mixing ratios for the 8-12 km region

for all latitudes were 0.4-0.6 ppt, with the highest concentrations observed again in the equatorial latitudes.

Halomethanes CH2BrCl, CHBrCl2, and CHBr2Cl are

thought to be produced by similar mechanisms and therefore have similar source distributions (Carpenter et al.,

1999). CH2BrCl has the longest average lifetime (Table

2-4) of the three gases and thus the highest overall concentrations. Schauffler et al. (1999) reported CH2BrCl

mixing ratios of 0.15-0.25 ppt at 10 km in the tropics,

decreasing to about 0.01 ppt at 20 km. CHBrCl2 concentrations were 0.1-0.15 ppt at 10 km and dropped to below

0.03 ppt at 15 km. Tropical concentrations of CHBr2Cl

were 0.05-0.2 ppt at 10 km and decreased to below 0.02

ppt above 15 km. These observations are consistent with

photolysis being a significant removal process for

CHBr2Cl and CHBrCl2 (see Section 2.3).

2.4.1.3 IODINE COMPOUNDS

Iodocarbons are mainly emitted from marine

sources, and their generally very short lifetimes are

reflected in their concentrations, typically 0.01-1 ppt.

Methyl iodide (CH3I) is the most prevalent iodocarbon in

the atmosphere. CH3I concentrations in the MBL are

higher at low latitudes and midlatitudes, and show clear

seasonal variations at midlatitudes with the highest concentrations in summer (Blake et al., 1999a; Yokouchi et

al., 2001). Continental CH3I concentration measurements

are scarce. Bell et al. (2002) simulated the distribution of

CH3I in the troposphere using a 3-D model and an ocean

source parameterized in terms of ocean productivity. The

model calculated a mean atmospheric lifetime (defined as

atmospheric burden divided by photochemical loss) of

CH3I against photolysis of about 6 days. The ratio of

upper tropospheric (8-12 km) to lower tropospheric (0-2

km) concentrations in the tropical marine environment

ranges from 0.11 in subsiding regions to 0.40 in upwelling

regions (Bell et al., 2002). In the upper troposphere and

lower stratosphere, CH3I has been observed at concentrations of 0.0-0.2 ppt (Blake et al., 1999a; Sturges et al.,

2001). Bell et al. (2002) calculated a net sea-to-air flux

of 210 Gg yr–1, which falls within the range of previous

estimates (100-500 Gg yr–1). Photochemical production

of CH3I in the seawater is consistent with the recent MBL

observations (Yokouchi et al., 2001). Biomass burning

(<10 Gg yr–1; Andreae et al., 1996), rice fields (90 Gg yr–1;

Redeker et al., 2000; Muramatsu and Yoshida, 1995), wetlands (7.3 Gg yr–1), and peatlands (1.4 Gg yr–1; Dimmer

et al., 2001) are the main continental sources. However,

higher emissions from biomass burning of 20 Gg yr–1 have

also been reported (Muramatsu and Yoshida, 1995).

Carpenter et al. (2000) reported the automated

measurements of biogenic alkyl halides at Mace Head,

Ireland, including the detection of CH2BrI and the first

non-Arctic measurements of CH2I2 and CH2ClI. The

atmospheric mixing ratios of CH2I2 (<0.02-0.36 ppt),

CH2ClI (<0.02-0.21 ppt), CH2BrI (<0.02-0.32 ppt), C2H5I

(<0.02-0.21 ppt), and CH3I (0.12-1.47 ppt) were used to

calculate iodine atom emissions. The largest contribution

of iodine atoms to the marine boundary layer was from

CH2I2 (lifetime on the order of a few minutes). Annual

global atmospheric input of 0.04 Gg yr–1 by macroalgae

was estimated by Giese et al. (1999). Thus, polyhalogenated iodine compounds released from macroalgae

make a very small contribution to the global iodine budget

(only 0.01-0.05% of total CH3I emission), but are likely a

significantly source of iodine to coastal atmospheres

(Bassford et al., 1999; Baker et al., 2000).

VERY SHORT-LIVED SUBSTANCES

2.29
Current estimates of total inorganic iodine concentrations in the stratosphere are based on column measurements of IO (Wennberg et al., 1997; Pundt et al., 1998;

Wittrock et al., 2000). The estimated range is from less

than 0.05 ppt to 0.7 ppt.

2.4.2 Estimates of Contributions from VSL

Halogen Source Gases

In this section, we describe a simple method for

estimating the percentage of halogen atoms (emitted as

organic source gases at the surface) that are transported to

the stratosphere via the two pathways identified in Figure

2-1. Numerical estimates can be obtained using observed

concentrations of the source gases and their degradation

products, or using model output. The paucity of observations limits this analysis to providing an order-ofmagnitude estimate. The exercise also helps to identify

useful diagnostics for comparing model results. For simplicity, we assume that the intermediate products are

sufficiently short lived that their contributions to the

PGI pathway are small.

Consider a source gas X that contains halogen

atoms and is emitted at E moles (hX) yr–1, where hX is inorganic halogen species (Cl, Br, or I) found in compound X.

The fractions of the emissions that enter the stratosphere

via the SGI pathway and the PGI pathway 

are given by

(2.1)

where and are the mixing ratios for X and hX near

the tropical tropopause, w is the large-scale upward velocity across the tropopause, MTT is the air density at the

tropical tropopause in moles per unit volume, and A is the

area of the Earth. To get a numerical value for , we

note that 

(2.2)

where is the replacement time for tropospheric air

against transport out of the troposphere, z is altitude, H0 is

the atmospheric scale height of air, Msurf is the air density

(in moles per unit volume) at the surface, and ZT is the

altitude of the tropopause. Rosenlof and Holton (1993)

provided an estimate of the mass flux across 100 mb that

can be used to give a value of 17 years for .

The emission E is equal to the removal rate of X in

the stratosphere and troposphere. We derive an expression for the removal rate assuming that the removal in the

stratosphere can be ignored and that the local lifetime is

uniform in the troposphere:

(2.3)

where , , and are the air density,

mixing ratio, and local lifetime of X as a function of .

The local photochemical lifetime of X is taken from

Table 2-4 and assumed to be approximately uniform in

the troposphere, is the mixing ratio of X in the boundary layer, and (z) is the normalized altitude profile of

and has the value 1 in the 1-km-thick BL and falls

off with the assumed scale height given in Table 2-10.

Note that although (z) is used to estimate the burden

of the source gas, is expected to be larger than 

because convection will carry additional source gas to the

TTL. Model results in Figure 2-4 clearly show that the

mixing ratio between 500 and 200 hPa do not follow a

simple exponential decrease with height. The contribution to the burden from the extra source gas carried by

convection to the TTL is ignored in this estimate.

Combining Equations (2.1), (2.2), and (2.3) gives

(2.4)

where is a simple quadrature in terms of the

scale heights (see footnote b in Table 2-10).

Table 2-8 gives the observed median surface concentrations and the median concentrations at 10 km in the

tropics for several VSL source gases from the TRACE-P

and PEM-Tropics A and B campaigns. In lieu of values

for , it is assumed that X is well mixed in the TTL and

~ . Using the values in Equation (2.4)

gives estimated values of (see Table 2-10) ranging

from 0.03% to 3% for source gases with local lifetimes of

5 days to 150 days. The values derived for the VSL source

gases with 100-day lifetimes are comparable with previous

FSGI

X

f f X X

UT BL f / X

UT f X

TTLf X

TTKH H f

X (, ) 0

FSGI

TT

BL

local

res trop

air

X

X

X

X

f

X f

f =¥ ¥ KH H

-

t

t ( , ) 0

f zH X

T f

X

BL ¥ -- exp[ ] ( )1 / f X

TTf X

norm

H f

X

f x X ( ) r f X

normf X

BL

r

x

t local

X ( ) x

r f x X ( ) r M( ) x

r

E = n Mx f x

x dV

n M f A e f z dz

X

h

X

X

X

h X

z

H X

Z

X

X

X

T

¥ ¥

a

¥ ¥ ¥¥ ¥

ÚÚÚ

Ú

-

() ()

( )

( )

r r

r t

t

local

surf BL norm

local

0

t res trop

air

-

t res trop

air

-

wM

M e

z

H

Z

dz

T

TT

surf

res trop

air =

¥ -

Ú

-

0

0

t

wMTT f hX

TT f X

TT

F E

F E F

SGI

TT TT

PGI

TT TT TT

TT

SG

X X

h X

X

h h

X

h X

X

n wM A

wM A f

n

X

X X

X

f

f

f

= ¥¥ ¥

= ¥ ¥ = ¥ ¥

( ) FPGI

X ( ) FSGI

X

nX

hX

VERY SHORT-LIVED SUBSTANCES

2.30
VERY SHORT-LIVED SUBSTANCES

2.31

Table 2-10. Values used to estimate the fraction of halogen atoms transported to the stratosphere via

the SGI pathway ( ) (see Section 2.4.2).

C2HCl3 4.6 1 4.8 0.2 0.0008

N/A N/A

N/A N/A

CH3I 7 1 4.8 0.05 0.00026

0.2 0.0011

0.15 0.000803

CHBr3 26 2 3.3 0.3 0.0042

0.3 0.0042

0.2 0.0028

CH2ClCH2Cl 70 5 2.0 0.3 0.0065

0.6 0.013

N/A N/A

CHBr2Cl 69 5 2.0 0.6 0.013

0.5 0.011

0.5 0.011

CHBrCl2 78 10 1.5 0.7 0.013

0.6 0.011

0.5 0.0093

C2Cl4 99 10 1.5 0.3 0.0071 0.01 *1

0.6 0.014

1.1 0.032

CH2Br2 120 15 1.3 0.8 0.020

0.7 0.018

0.8 0.020

CH2Cl2 140 15 1.3 0.5 0.015 0.014 *2,

0.7 0.021 0.02 *3

1.1 0.033

CHCl3 150 15 1.3 0.6 0.019 0.004 *4, 

0.8 0.025 0.02 *1

1.0 0.031

In columns 5 and 6, for each species, the first value is for data from TRACE-P; the second, from PEM-Tropics B; and the third, from PEM-Tropics A.

References: *1 Kindler et al. (1995); *2 Keene et al. (1999); *3 Graedel and Keene (1995); *4 Khalil (1998), results are from 12-box model.

a From Table 2-4. b

.

The integral is calculated assuming that areas in the tropics and extratropics are equal and ZT is 15 km and 10 km for the tropics and the extratropics,

respectively. See Section 2.4.2 for definitions of other terms. c From Table 2-8.

K H H e dz e f z dz f z z f z e z f

X

z

H

Z z

H

Z

XX X

z

H T T

f

X

( , ) () () () 0

0 0

1

0 0 = × =≤ = ≥ 11 1 − − − −

∫ ∫ normal normal normal where for ; for km km

FSGI

X

Species Local Estimated b Estimated : c

Lifetime a

(days) (km) TRACE-P

PEM-Tropics B

PEM-Tropics A

H f

X

f

f

X

X

UT

BL

KH H f

X (, ) 0 F Calculated SGI

X Model

Using Eq. (2.4): Estimate of

TRACE-P

PEM-Tropics B

PEM-Tropics A

FSGI

X
results derived using a 12-box numerical model (see reference cited in Table 2-10).

For , we use the following estimate:

(2.5)

where tPG-X is the local residence time (against chemical

degradation, transport, and washout) for the product gas

from the degradation of X. If there is no long-lived intermediate product for species X, tPG-X will correspond to the

washout lifetime of the inorganic halogen species. Using

a nominal value of 10 days would indicate that is comparable with for source gases having local lifetimes

of about 10 days, and a tenth of for source gases

having lifetimes of order 100 days. If there is a long-lived

intermediate (tPG-X longer than half a year), the value for

could approach (1- ).

It is important to recognize the assumptions that

were used to obtain these estimates. In the procedure, we

assumed that the flux through the tropopause is governed

by the large-scale circulation with uniform and 

along the tropopause and that the concentrations of X or

hX are uniform within the TTL so that observations collected in the upper troposphere can be used. In addition,

we did not include transport along isentropic surfaces

in the subtropics or sub-synoptic-scale stratospheretroposphere exchange events that occur in the extratropics.

In estimating the chemical removal rate of the source gas,

it is assumed that the local photochemical lifetime is uniform in the troposphere. A similar assumption is made

for the washout rate of the inorganic halogen. However,

these two parameters are likely to vary strongly with latitude and season and to be very dependent on the location

of emissions.

To relate the concentration of inorganic halogen in

the stratosphere to the concentration of the source

gas at the surface , one can proceed as follows. For

simplicity, it is assumed that the source gas contains one

halogen atom. Take the case in which the concentration

of the source gas is 1 ppt in a 1-km layer above the surface and the lifetime of the gas is 10 days. The emission

needed to balance the removal in the BL is estimated to

be 8 ¥ 108 moles yr–1. If one assumes that 0.1 to 1% of

the flux emitted in the troposphere enters the stratosphere

as odd halogen, the flux is 0.8 ¥ 106 to 8 ¥ 106 moles yr–1.

Taking a stratospheric residence time of 2 years (see, e.g.,

Holton, 1990) would mean that the average mixing ratio

for inorganic halogen in the stratosphere is about 0.1 to 1

ppt. The typical concentrations for VSL halogen source

gases are of order several ppt for bromine and iodine

source gases, and up to 50 ppt for chlorine source gases.

The current inorganic chlorine concentration in the stratosphere is about 3 ppb. It is thus unlikely that VSL chlorine source gases will make a significant contribution to

the chlorine budget in the middle and upper stratosphere.

However, they may still make a contribution close to the

tropopause. In the case of bromine (with inorganic bromine concentration at 20 ppt and iodine at <1 ppt), it is

likely that the effects from VSL halogen source gases will

be important throughout the stratosphere.

2.4.3 Modeling Studies

We now review some of the modeling work on the

contribution of VSL source gases to the stratospheric inorganic halogen budget. There have been two recent model

studies on bromoform (CHBr3). In both studies it was

assumed that the degradation products immediately

released their Br atoms. Dvortsov et al. (1999) used a

global two-dimensional (2-D) model with parameterized

convection and washout prescribed as a function of altitude, resulting in washout time constants ranging from

about 9 days at the ground to 17 days just below the

tropopause. Emissions were assumed to be uniform over

the ocean and were adjusted to give a surface concentration of 1.5-2 ppt CHBr3. In their model simulation, sufficient bromine atoms (in its organic and inorganic form)

were transported to the stratosphere to maintain ~1 ppt

Bry. Note that 75% of the flux came through the tropical

tropopause as Bry. Nielsen and Douglass (2001) ran an

offline 3-D model using assimilated winds to simulate

CHBr3 and Bry. Their washout time constant was fixed at

~6 days throughout the troposphere. Surface fluxes and

mixing ratios were treated in a manner similar to that of

Dvortsov et al. (1999). Figure 2-9 shows their average

modeled tropical CHBr3 profile again compared with

PEM-Tropics data, plus measurements in the tropical

tropopause layer (TTL) from the STRAT campaign.

Nielsen and Douglass (2001) also calculated that ~1 ppt

Bry is derived from CHBr3 in the lower stratosphere, with

50% of this amount transported as inorganic bromine.

Although both model studies agree well with data

up to about 10 km, both studies overestimate CHBr3 in

the TTL compared with the STRAT data. Because the

models maintain too much CHBr3 in the TTL, Nielsen and

Douglass (2001) estimate that CHBr3 contributes to maintaining, at most, 15% of the Bry in the stratosphere. The

relative effect of this contribution is likely to be greatest

in the layer immediately above the tropopause, where Bry

from CHBr3 would increase total Bry by ~45%. These

( f ) X

BL

( ) f hX

Strat f hX

TT f X

TT FSGI

X FPGI

X

FSGI

X

FSGI

X

FPGI

X

F FF PGI

TT

TT

SGI

PG

local

SGI

X

h

X

h X

X X

X

X f

n f

X

X

= ¥ ¥a ¥ - t

t

FPGI

X

VERY SHORT-LIVED SUBSTANCES

2.32
modeling studies, together with the observationally based

studies of Wamsley et al. (1998) and Pfeilsticker et al.

(2000), suggest the importance of including shorter lived

species in the stratospheric bromine budget.

Several recent studies provide estimates for the

impact of 1-bromopropane (n-PB) on stratospheric

bromine levels (Olsen et al., 2000; Bridgeman et al., 2000;

Wuebbles et al. 2001) (see Section 2.5.3). They all noted

that the idea of a single lifetime is flawed for VSL species

and that mixing ratios of n-PB reaching the tropical upper

troposphere depend strongly on the latitude and season of

the emissions. The Wuebbles et al. (2001) study used a

global 3-D model with a parameterized chemistry scheme

to investigate transport of n-PB and its degradation

products across the tropical tropopause. They found that

tropical emissions were twice as likely to enter the stratosphere as high-latitude emissions, in qualitative agreement with the other two studies. Wuebbles et al. (2001)

showed that up to 0.5% of the Br emitted as n-PB at the

surface entered the stratosphere, with only one-third of Br

in the form of n-PB (the SGI pathway). The remainder

was mostly in the form of inorganic bromine (the PGI

pathway). However, these values are likely to change as

more reliable information on the degradation of n-PB

becomes available.

All the modeling studies emphasize that the lack of

knowledge about the fate of the degradation products

limits our ability to quantify their contribution to halogen

loading in the stratosphere. Up to now, little attention has

been paid to evaluation of model simulations of the

processes and time scales necessary to understand the

transport and chemistry of VSL substances. For example,

cross-tropopause transport rates are very important for

determining the effects of Bry on O3 loss in the lowermost

stratosphere, yet neither observations nor models are able

to fully agree on the degree of communication between

the troposphere and stratosphere (see Section 2.2.2.2).

Furthermore, few models had their transport evaluated in

this region against observations. Examples of such studies

include the analyses by Strahan et al. (1998) and Douglass

et al. (1999). These studies evaluated the tropopause

behavior of different models participating in the Global

Modeling Initiative by comparisons with the known phase

lag between UT and LS CO2. Clearly more studies of this

type are required.

2.4.4 Information Needed for Future

Evaluation

In summary, it can be concluded that the effects of

VSL halogen source gases on halogen loading in the stratosphere are likely to be limited to bromine and iodine

budgets in the present-day atmosphere. It is unlikely that

VSL chlorine compounds will make a significant contribution to the chlorine budget.

It is clear that more observations of VSL substances, their degradation products, and inorganic

halogen species are required in the troposphere and the

stratosphere if we are to reduce the uncertainties surrounding our estimates of the contribution of these compounds to the present halogen loading in the stratosphere.

In particular, measurements are needed in the TTL and

the lowermost stratosphere. This could be particularly

challenging if the sequestration of halogen ions in aerosol

particles turns out to be important (Murphy and Thomson,

2000). It is encouraging that several independent models

simulate similar amounts (~0.5%) of Bry input into the

stratosphere from n-PB, although this cannot be easily

verified using available observations. The observed

ratios as defined in Table 2-8 for specific source

gases could provide useful constraints although one

must also have a good knowledge of the distribution of

surface source strengths. Improved estimates will be

brought about by improvements in modeling capabilities

f f X X

UT BL /

VERY SHORT-LIVED SUBSTANCES

2.33

0.0 0.4 0.8 1.2 1.6 2.0

pptv

0

5

10

15

20 Altitude (km)

PCTM

PEM-Tropics

STRAT

320

340 Potential temperature (K)

400

380

360

Figure 2-9. Results from 3-D model simulations

(PCTM, parameterized chemistry and transport

model) of tropical CHBr3 and comparison with PEMTropics and STRAT data within 30° of the equator.

The different symbols denote data taken at different

times (diamonds for data taken in January and

February, triangles for July and August, and squares

for December). The model overestimates the transport of CHBr3 to the tropical upper troposphere (see

Nielsen and Douglass, 2001). Figure courtesy of A.

Douglass, NASA GSFC, U.S.
and parameterization of processes like convection and

washout. Cross-tropopause transport in the tropics,

through the TTL region, also requires thorough evaluation in global models including tests of tropical ascent

rates and exchange with the extratropics. These studies

need to be coupled to better understanding of degradation products, their local lifetimes and solubilities (also

see Section 2.3).

2.5 ESTIMATES FOR THE POTENTIAL IMPACT

OF HALOGENATED VSL SUBSTANCES ON

COLUMN OZONE

An operational definition for the Ozone Depletion

Potential (ODP) of a species X is given by

(2.6)

where DO3(X, E) is the steady-state percent change in

stratospheric ozone burden due to emission of X at the rate

of E g yr–1 at a specific location ( ) and time (te) of the

year. The ODP value is defined relative to the ozonedepleting efficiency of CFC-11 (CCl3F). An equivalent

definition for ODP is the cumulative ozone depletion for

the same mass pulse emission (same location and time of

year) of X versus that of CFC-11 (Prather, 1996, 1997,

2002). This latter definition leads naturally to using the

ODP values in equivalent effective stratospheric chlorine

(EESC) calculations (Daniel et al., 1995; Madronich and

Velders et al., 1999).

Figure 2-10 shows the information needed and the

procedure for calculating the change in column ozone

from emission of a halogen source gas at the Earth’s surface. The procedure for calculating ODP was originally

developed for halocarbons (Wuebbles, 1981, 1983) with

lifetimes sufficiently long (more than 1 year) that their

mixing ratios are uniform in the troposphere and the

increase in inorganic halogen in the troposphere is

expected to be small. For those gases, the processes in

the yellow boxes in Figure 2-10 are unimportant, and thus

the ODP value is independent of location and time of emission. The methods (either using results from numerical

models or using measured quantities in semiempirical

approaches) for calculating ODP values for long-lived

source gases have been reviewed in previous reports, and

they are judged to give reliable results. However, many

of the processes in the yellow boxes and how well they

are simulated in models have not been evaluated. Thus,

assessing the reliability of the ODP values for very shortlived (VSL) halogen source gases emitted at the Earth’s

surface requires further evaluations.

r

xe

ODP O

O CFC (, ,) (,)( ,) Xx t X EE e e

r = -

D

D

3

3 11

VERY SHORT-LIVED SUBSTANCES

2.34

Distributions of

inorganic halogen & 

intermediate products

in the stratosphere

Kinetic data 

for source gas

Ozone chemistry

in the stratosphere

Unimportant for long-lived source gases

Transport in

 troposphere

Transport in

 troposphere

Distribution of

source gas in the

troposphere

Distributions of

inorganic halogen &

intermediate products

in the troposphere

Ozone chemistry

in the troposphere

 Stratospheric

 ozone

response

SGI pathway

PGI pathway

Source gas

Intermediate and

Final products

Transport in

stratosphere

Tropospheric

 ozone

response

Emission

location

Kinetic data for

degradation

products

Emission

mass

Strat/trop

exchange

Figure 2-10. Schematic showing information needed to assess the ozone depletion estimate of halogen source

gases. Items in rectangles are species specific. Items in ovals are common to all species. Items in parallelograms are model predictions. Several processes are highlighted in two yellow-colored boxes with heavy dashed

outlines. The simulation of these processes in numerical models has not been fully evaluated.
Several recent efforts to address these issues

include a workshop on VSL compounds in March 1999

sponsored by the U.S. Environmental Protection Agency

(EPA) and the National Aeronautics and Space

Administration (NASA) (Wuebbles and Ko, 1999); the

findings of a special United Nations Environment

Programme (UNEP) report (UNEP, 2000); and several

recent publications (Bridgeman et al., 2000; Olsen et al.,

2000; Wuebbles et al., 2001). We review the methods

used to compute ODP for long-lived source gases in

Section 2.5.1, and in Section 2.5.2 we contrast them with

the new method needed for the VSL source gases. Some

of the results for ODP based on numerical models are

reviewed in Section 2.5.3. A brief discussion on performing the chlorine loading calculation is presented in

Section 2.5.4.

2.5.1 Review of Methodology for

Calculating ODPs for Long-Lived

Source Gases

Because the ODP value of a long-lived source gas

is independent of the location and time of emission, we

will suppress the and te notations in this section.

Solomon et al. (1992) provided an expression for computing a semiempirical value for the ODP (ODPs-e(X)) of

a long-lived source gas. This is useful for relating ODP

values to measured quantities that could be used to evaluate the model-calculated results. We will follow a similar procedure here and express DO3(X, E) as a product of

two factors:

DO3(X, E) = HLX(E) ¥ dX (2.7)

where HLX(E) is the steady-state stratospheric inorganic

halogen loading (chlorine, bromine, or iodine mixing ratio

in mole fraction) associated with emission of X at E g yr–1

and dX is the ozone-removal efficiency per unit inorganic

halogen loading from X. Note that the definitions for

HLX(E) and dX are related in that there is more than one

way to quantify the halogen loading. One could either

use the average mixing ratio of the inorganic halogen

in the stratosphere or the mixing ratio at the

stratopause , where Strp means stratopause).

Depending on which definition one chooses, dX will take

different values. Using this notation,

(2.8)

where CLP(X) is the chlorine loading potential and VX is

the ratio of the ozone-removal efficiency factors. For

long-lived source gases where the contribution from the

PGI pathway to the inorganic halogen loading in the

stratosphere is small compared with that from the SGI

pathway, the steady-state mixing ratio of hX at the

stratopause (that results from a constant emission of X) is

approximately equal to the mixing ratio of X in the troposphere (see Prather and Watson, 1990). Now

(2.9)

where mX is the gram molecular weight, tX is the atmospheric lifetime of X, is the number of halogen atoms

in X, and M is the number of moles of air in the atmosphere. Thus, for long-lived species like those treated in

Chapter 1,

(2.10)

The efficiency factor has two components:

. The values for the first component (VS

, S

for speciation) depend on whether the halogen is chlorine,

bromine, or iodine. The other factor ( , D for distribution) has to do with the distribution of the inorganic

halogen in the stratosphere. Equation (2.8) corresponds

to the same definition used in Chapter 1, with VS identified with the a factor, and is the fractional release

factor. When a numerical model is used to calculate

ODP(X) via Equation (2.6), one can identify the value for

with the ODP(X)/CLPStrp(X) ratio. From this perspective, the value for corresponds to a spatial

and seasonal average weighted by the ozone response in

the lower stratosphere. Thus, it is not possible to pick

values for VS and at a specific location and time to represent the ODP(X)/CLPStrp(X) ratio. A way to proceed is

to pick a conventional value for V S and define as

ODP(X)/CLPStrp(X) divided by VS

. The value for VCl is 1

by definition. Based on the ODP(X)/CLPStrp(X) ratios for

chlorine species (see Table 1-5), the value for ranges

from 0.35 for HCFC-22 (CHClF2) to 1.08 for methyl chloroform (CH3CCl3). Model simulations by Ko et al. (1998)

gave a value of 58 for for CH V V 3Br. The con- Br

CH Br 3 ¥ D

V X

D

V X

D

V X

D

V V S

X

D ¥

V V S

X

D ¥

V X

D

V X

D

V V V X

S

X

Strp D = ¥ V X

Strp CLP HL

HL

Strp

Strp

CFC

Strp

CFC

CFC CFC

Cl

( ) ( )

( ) X E

E

n

n

X

X

X X

hX

=

= ¥¥

-

-

- -

11

11

11 11

m

m

t

t

nX

hX

HLStrp

X

X

X

X

h E E n X ( ) = ¥ ¥ ¥ 1

M

t

m

ODP HL

HL

CLP

CFC CFC

s e

X X

X

X E

E

X

-

- -

= ¥

= ¥

( ) ( )

( )

( )

11 11

d

d

V

( ( )) HLStrp

X E

( ( )) HLAvg

X E

r

xe

VERY SHORT-LIVED SUBSTANCES

2.35
vention is to take ςBr to be 45, implying that has the

value of 1.3. The kinetic data are less complete for iodine.

Solomon et al. (1994a) provided an estimate of 100-2000

for ςI based on estimated reaction rate constants. Bedjanian et al. (1998) gave 150 using updated rate data. A

more recent study using JPL-2000 rates (Sander et al.,

2000) from theAtmospheric and Environmental Research,

Inc. (AER) model (Ko et al., 1998) gave a value of 300

for for CH3I. This could be interpreted as 150

for ςI and 2 for . However, considering the missing

data on some of the reservoir species (see Section 2.3.4),

we will use a range of 150-300 for ςI. A useful guideline

for is that is greater or smaller than 1 depending

on whether the local lifetime of X in the lower stratosphere

is shorter or longer than the local lifetime of CFC-11 (~2

years). From the values of fractional release measured

for CFC-11 in the lower stratosphere, should generally not be much larger than 2. Thus, the estimate for longlived source gases is

(2.11)

2.5.2 Modified Approach for Estimating

ODP for Halogenated VSL

Substances

The inorganic halogen loading from a VSL source

gas is maintained by the transport of halogen atoms

affected by the SGI and PGI pathways. Solomon et al.

(1994b) suggested a method to estimate the contribution

from the SGI pathway based on the observed concentration of the VSL source gas near the tropopause. We will

restrict the discussion in this subsection to emissions at

the Earth’s surface. Thus, can be represented by θe

and φe, corresponding to the latitude and longitude of the

emission location. It should be noted that certain applications (e.g., inerting fuel tanks in aircraft) involve release

of the chemical in the atmosphere above the surface. In

those cases, the ODP value will also depend on the altitude at which the VSL substance is released. In this section, using the notation from Section 2.4 (Equation (2.1))

and , we obtain

(2.12)

where MStrat is the number of moles of air in the stratosphere, and E, , and are functions of θe, φe, and

te. The residence time for the inorganic halogen in the

stratosphere is designated by . Strictly speaking,

could also be a function of θe, φe, and te because

material that enters the tropical lower stratosphere has residence times different from material that enters the extratropical lower stratosphere. We do not consider this for

now.

The expression for is much simpler:

(2.13)

where we used the facts that for CFC-11, ,

, and it contains three chlorine atoms. We left off

the explicit θe, φe, and te dependence because the loading

for CFC-11 will only depend on the averaged annual tropospheric emission independent of location.

Using Equations (2.12) and (2.13) in Equation

(2.8), we obtain

(2.14)

Similar to the discussion in Section 2.5.1, we use the

expression

where is the distribution factor for hX maintained

by either the SGI or PGI pathway. Thus, for VSL com-

 ̃

ζ X

D

−SGI/PGI

ς X ς ζ θ φ S

X

D

e e et −SGI/PGI =− × ( ) AvgSGI/PGI  ̃ , ,

ODP ODP

ODP

-

VSL

- -SGI

VSL

- -PGI

VSL

CFC- res-strat

res-strat

-SGI

Avg

SGI

-PGI

Avg

CFCs e e e e s e e e e

s e e e e

X

X

h h

h

X

X

e e e

X

X t X t

X t

n

t

X X

( , , , ) ( , , , )

( , , , )

( ( , , ))

(

θ φ θ φ

θ φ

μ

μ

τ

τ

ς θ φ

ς

=

+

= × ×

× { ×

+ ×

11

3 11

F

FPGI ) X

e e e (θ , φ , t ) }

FPGI

X = 0

FSGI

X =1

×

×

×

−

−

E h −3

11

11

μ

τ

CFC

res strat CFC 1

MStrat

HLCFC

Avg

−11(E) =

HLCFC

Avg

−11(E)

τ res−strat

hX

τ res−strat

hX

FPGI

X FSGI

X

HL HL

HL

Avg

-SGI

Avg

-PGI

Avg

Strat

res strat

SGI PGI

X e e e X e e e

X e e e

e e e X

h

X

h

X

e e e

X

e e e

E t E t

E t

E t n

t t

X

X

( ( , , )) ( ( , , ))

( ( , , ))

( , , )

( ( , , ) ( , , ))

θ φ θ φ

θ φ

θ φ

μ

τ

θ φ θ φ

=

+

= ×

×

×

× +

−

M

F F

E = E × nX

h

X X / μ

r

xe

ODPlong lived CFC

CFC

Cl

Br

I

s e

X

X X

h

X

D

X n X

−

− −

−

= × ×

×

=

=

≈ →





















×

( ) μ

μ

τ

τ

ς

ς

ς

ς

11

11 3

1

45

150 300

ς X

D

ς X

D ς X

D

ς CH I3

D

ς ς I

CH I3

× D

ς CH Br 3

D

VERY SHORT-LIVED SUBSTANCES

2.36

− −
pounds, the concept of a single ODP generally provided

to policymakers for long-lived species needs to be

replaced with a table of ODP values, essentially a matrix

of ODP values depending on the location and season of

emissions. To obtain an order-of-magnitude estimate for the

ODP values, we assume has values between

0.001 and 0.01 for a VSL source gas with no long-lived

intermediate products. We will also assume that the

product is of order 1. This is

reasonable because a species with smaller fractional

release in the lower stratosphere will have a larger .

These values give

(2.15)

where the notation VSL* serves as a reminder that the estimate is valid only for VSL source gases with no long-lived

intermediate products.

The ODP estimate given by Equation (2.15) represents ozone depletion in the stratosphere. If degradation

of the source gas leads to sufficiently high local concentrations of inorganic halogen species in the free troposphere, there can be a decrease in ozone in the free troposphere on a regional scale. Observed depletion of ozone

in the boundary layer of the polar region following polar

sunrises has been attributed to inorganic bromine (see,

e.g., Barrie and Platt, 1997). In polluted regions, the VSL

halocarbons can also act as a volatile organic compound

(VOC) to produce ozone. The expected tropospheric

ozone response will depend on the NOx, HOx, and hydrocarbon distributions in the region. Thus, the expected

response will again depend on the location and time of

emission. This raises an interesting issue as to whether

and how one should include such effects in the ODP

estimates.

2.5.3 Existing Modeling Studies on ODP

for n-PB

If one is confident that a numerical model can correctly predict the halogen loading and ozone depletion in

the stratosphere, one can then use the model results in

Equation (2.6) to compute an ODP. Two-dimensional

models of the stratosphere have been the accepted tools

for calculating the ozone response in the stratosphere.

Most of them lack parameterizations for convection and

do not represent realistically washout processesin deriving

the transport of the VSL source gases and their degradation products to the stratosphere. Three-dimensional

chemical transport models (CTMs) are more suited for

studies of tropospheric processes but generally do not have

complete representations of the stratosphere and its

processes. In both cases, the model results must be interpreted with care because there has been no study that comprehensively evaluates how well models simulate the

processes in the troposphere.

Many of the lessons learned come from studies of

n-propyl bromide. Earlier studies (Nelson et al., 1997;

Wuebbles et al., 1998, 1999) based on analyses assuming

emissions evenly distributed across the globe had derived

an atmospheric lifetime for n-PB due to its reaction with

OH of 11-14 days. The single-valued ODPs for n-PB

using the traditional approach in Wuebbles et al. (1998,

1999) have been supplanted by the following analyses

using the new approach.

Olsen et al. (2000) applied the low-resolution (8°

latitude × 10° longitude × 9 layers) three-dimensional

University of California at Irvine (UCI) tropospheric

chemical transport model to determine the amount of VSL

substances reaching the stratosphere via the SGI pathway

for a range of atmospheric lifetimes, and for various

seasons and latitudes of emission. The model-calculated

inorganic halogen loadings from n-PB sustained by the

SGI pathway were used as a scaling parameter to obtain

ODP values for different emission scenarios based on the

relation between ODP and loading computed from a 2-D

model (Wuebbles et al., 1999). The estimated ODP values

from Olsen et al. (2000) range from 0.0002 (for summer

emissions at 56° to 64°N) to as large as 0.06 for tropical

emissions (as shown in Table 2-11).

Bridgeman et al. (2000) used the TOMCAT threedimensional global chemical-transport model (5.6° ×

5.6° × 31 levels) to evaluate the atmospheric lifetime and

amount of n-PB reaching the stratosphere as a function of

location and season of emission. They then used an

empirical approach to the calculation of ODPs, based on

the derived lifetime and the fraction of n-PB emissions

reaching the stratosphere. For n-PB, they derived ODPs

that vary from 0.0033 for emissions from Europe to

0.0109 for emissions from Indonesia. The study assumed

that any bromine atoms released in the troposphere would

be removed by rainout with 100% efficiency before

reaching the stratosphere. Thus, the contribution from

the PGI pathway was ignored. They did suggest that such

transport of degradation products could significantly

increase the amount of bromine reaching the stratosphere.

Wuebbles et al. (2001) used the Model of Ozone

and Related Chemical Tracers, version 2 (MOZART2)

ODPVSL* CFC

Cl

Br

I

s e

X

X

h n X

−

−

≈ → ×

× ×

=

=

≈ →





















(0.001 0.01)

3

1

45

150 300

μ 11

μ

ς

ς

ς

τ res−strat

hX

τ τ ζ res strat res strat SGI/PGI CFC

− − −

−

( ) × h h

X X D / 11  ̃

FSGI FPGI

X X +

VERY SHORT-LIVED SUBSTANCES

2.37
three-dimensional chemical-transport model (5° ¥ 5° ¥

34 levels) in combination with their two-dimensional

model to compute ODPs for n-PB. An effort was made to

quantify the effects of degradation products. With thenavailable information, bromoacetone was considered to

be the intermediate degradation product with the largest

concentration, having a local lifetime of about 1 day. The

three-dimensional model was used with simplified chemistry having bromoacetone as the only intermediate degradation product. For n-PB emitted uniformly over global

land masses north of 60°S and south of 70°N, their threedimensional model calculations suggest that about 0.44%

of the emitted bromine enters the stratosphere. About 33%

of the bromine reaching the stratosphere results from the

direct transport of n-PB to the stratosphere and about 19%

due to the transport of bromoacetone to the stratosphere.

The rest (48%) is from transport of inorganic bromine.

Calculated ODP values for other emission scenarios are

given in Table 2-12.

Recent data (Burkholder et al., 2002) suggest that

the lifetime of bromoacetone is of the order of hours rather

than days. Thus, the contribution from bromoacetone may

be considerably smaller in evaluating the ODPs for n-PB.

As an estimate of the effects of the newly evaluated much

shorter atmospheric lifetime, Table 2-12 uses the model

results in the Wuebbles et al. (2001) paper to estimate the

resulting ODPs if no bromoacetone had reached the stratosphere. The resulting ODP values are reduced, but only

slightly. However, the estimate did not account for the

likely increase in the amount of inorganic bromine that

would reach the stratosphere if the bromoacetone lifetime

was reduced to a few hours, and the different distribution

in inorganic bromine in the two cases.

In summary, the existing studies for n-PB have

shown that the ODP for a VSL source gas does vary with

location and season of the emissions, by more than a factor

of 5 for this particular gas, with the largest values for

tropical emission. However, there remain significant

uncertainties in the existing studies. Better treatments of

tropospheric convection processes and wet scavenging

processes could be particularly important to reducing the

uncertainties for ODPs.

2.5.4 Information Required for Evaluating

Ozone Impact from VSL Source

Gases

As discussed above, unlike for the longer lived

compounds, the ODP for the VSL source gases cannot be

viewed as an intrinsic property of the substance that is

equally applicable worldwide. The location and timing

of the emissions are important to determine a meaningful

set of ODP values for VSL halogen source gases. Figure

2-10 shows the different steps needed in estimating the

ozone depletion effects of long-lived species and VSL

halogen source gases. Future studies of ODPs for VSL

substances require careful consideration of the following

parameters:

VERY SHORT-LIVED SUBSTANCES

2.38

Table 2-12. Estimated ODP for n-PB as a function of location of emissions based on Wuebbles

et al. (2001).

Emission Derived Modified ODP

Scenario ODP Range

Range (Reduced

Bromoacetone)

Global 0.033-0.040 0.027-0.038

North America 0.018-0.019 0.015-0.017

United States 0.016-0.019 0.013-0.018

North America + 0.021-0.028 0.017-0.026

Europe + Asia

Indonesia + 0.087-0.105 0.071-0.100

Southeast Asia

Estimates are based on combined results from analyses with the

University of Illinois at Urbana-Champaign (UIUC) two-dimensional

model and the MOZART2 three-dimensional model. Emissions were

evenly distributed over the land masses based on their representation

in the three-dimensional model. The range represents different

assumptions on the degradation chemistry. In the simulation, bromoacetone is the major intermediate degradation product that is transported to the stratosphere. The last column gives the estimated ODPs

assuming that bromoacetone does not reach the stratosphere.

Table 2-11. Summary of existing studies that

have determined a range of ODP values for the

case of n-propyl bromide (n-PB), updated from

the evaluation in UNEP (2000).

Study Range of Effects of 

ODP Halogen

Values a Loading

from the

PGI

Pathway

Bridgeman et al. (2000) 0.0033-0.0109 No

Olsen et al. (2000) 0.0002-0.06 No

Wuebbles et al. (2001) 0.016-0.105 Yes

a Range for the different set of calculations performed in each study.

The extreme range of values in Olsen et al. (2000) was due to calculating separate results for individual seasons.
• Determination that the vapor pressure of the substance, at atmospheric temperatures at the point of

emission, is sufficient to generate a significant gasphase concentration in the atmosphere.

• Determination of the atmospheric reactivity of the

VSL substance (with OH, through photolysis, etc.)

and its water solubility.

• Evaluation of the degradation products, their water

solubility, and their chemical behavior in the troposphere.

• Determination of the potential uses of the VSL substance, where the VSL substance will be used and

likely emitted, and the likely seasonal variations in

the emissions (see, e.g., data compiled for n-PB by

the UNEP Technology and Economic Assessment

Policy (TEAP) task force (UNEP, 2001)).

Given the complexity of the processes in the troposphere,

it is unlikely that they could be treated in 2-D models.

Three-dimensional models of the global troposphere

and stratosphere are necessary to determine accurately

the resulting set of ODPs for a newly considered VSL

substance.

As discussed in Section 2.5.1, the ODP represents

the cumulated effects from the same mass pulse emission

relative to CFC-11. The time evolutions of the effects are

distinctly different for long-lived species and VSL source

gases. For long-lived species, the full effect is realized

after a few years’ delay and then decreases exponentially

according to the atmospheric lifetime of the source gas.

For CFC-11, this means that the cumulative effect is spread

over about 100 years. For VSL source gases, the effect

lasts about 5 years. Thus, the short-term (5 years after

emission) effects from a VSL source gas can be 20 times

larger than indicated by its ODP value. For fully analyzing

the potential effects on ozone, it is important to consider

the projected amount of future use and resulting emissions

of the VSL substances and evaluate the time-dependent

effects (e.g., Wuebbles and Calm, 1997; Calm et al., 1999).

One approach often used for evaluating future emissions

is the concept of equivalent chlorine loading, or more

appropriately for determining stratospheric ozone loss, its

variation, equivalent effective stratospheric chlorine

(EESC) (e.g., Solomon and Wuebbles et al., 1995; Daniel

et al., 1995; Madronich and Velders et al., 1999; Calm et

al.,1999). As demonstrated in Chapter 1, EESC can be

very useful for evaluating the effects on ozone from longlived halocarbons. However, given that the halogen

loading from VSL source gases as given by Equation

(2.13) is very different from Equation (2.8), one cannot

use the normal EESC formulation for VSL source gases.

A modified procedure is needed for such calculations.

2.6 VSL SULFUR SPECIES AND STRATOSPHERIC AEROSOLS

2.6.1 Introduction

Stratospheric aerosol consists mainly of supercooled hydrated H2SO4 particles with a sulfuric acid

mass fraction in the range of 50-80%. At very low temperatures in the wintertime polar stratosphere, nitric acid

can be taken up on sulfuric acid particles to form a ternary

solution HNO3/H2SO4/H2O. Stratospheric aerosols can

stay liquid to very low temperatures. Other aerosol components can also be present (Carslaw et al., 1997; Murphy

et al., 1998; Murphy and Thomson, 2000). The presence

of soot in the lower stratosphere has been observed and

documented (Pueschel et al., 1992, 1997; Strawa et al.,

1999), although measurements remain scarce. Smoke

from boreal forest fires has also been observed in some

occasions in the lower stratosphere (Fromm et al., 2000).

The particles form a persistent aerosol layer, the

stratospheric sulfate aerosol (SSA) layer or Junge layer,

which extends from the tropopause up to ~25 km altitude.

This layer shows strong variability at annual or longer

time scales due to the dominant influence of explosive

volcanic eruptions that inject sulfur compounds into the

stratosphere (see, e.g., Godin and Poole et al., 1999). For

example, satellite observations from the Stratospheric

Aerosol and Gas Experiment (SAGE) II show that it took

6 years for the stratospheric aerosol concentrations to

return to their pre-Pinatubo-eruption levels of 1991 (Kent

et al., 1995, 1998; Thomason et al., 1997a,b) after a 100-

fold increase in mass loading. It is generally believed that

the observed sulfate loading after the year 2000 is representative of the background sulfate aerosol unaffected by

large explosive volcanic eruptions.

Changes in stratospheric aerosols will affect stratospheric ozone in at least two ways. First, heterogeneous

reactions occurring on the aerosol surface affect the partitioning of chemical species between reservoirs and free

radicals. This is primarily due to the heterogeneous conversion of active NOx to NOy reservoirs, which significantly alters the partitioning of HOx and halogen species.

Thus, an increase in the aerosol surface area tends to convert more of the halogen reservoirs to radicals and hence

affects the catalytic cycles leading to ozone depletion.

Second, changes in aerosol loading change the radiative

balance, and therefore the temperature, in the lower stratosphere. This in turn changes the transport circulation,

resulting in a redistribution of ozone. Effects from both

mechanisms were evident after the Mt. Pinatubo eruption

(see, e.g., Pitari and Mancini, 2002).

VERY SHORT-LIVED SUBSTANCES

2.39
The purpose of this section is to examine the

processes that determine how injection of sulfur from

large explosive volcanic eruptions and release of VSL precursor gases in the troposphere might affect the SSA.

Stratospheric aerosol is maintained by the supply of sulfur

(in gas and liquid forms) to the stratosphere that is balanced with removal by transport out of the stratosphere.

The residence time for aerosols in the stratosphere is of

the order of 1 year. It is not possible to define a single

residence time for aerosols, because it changes with the

size distribution of the aerosol particles and it also depends

on where the aerosol particles are formed in the stratosphere. The supply to the stratosphere could be in the

form of sulfate (Brock et al., 1995) or gas-phase sulfate

precursors. Figure 2-11 summarizes the processes that

control the gas-phase sulfate precursor gases. Sulfate precursors released from the Earth’s surface include sulfur

dioxide (SO2), hydrogen sulfide (H2S), and organic gases

such as dimethyl sulfide (DMS, CH3SCH3), carbonyl sulfide (OCS), and carbon disulfide (CS2). DMS, H2S, and

SO2 can be oxidized in the lower troposphere via gasand/or aqueous-phase processes with a lifetime of a few

days or less, forming sulfate aerosol as the main final

product. Thus they behave like the VSL halogen source

gases discussed earlier in this chapter. The present emission rates have been estimated as 10-30 Tg (S) yr–1 for

DMS, 1-2 Tg (S) yr–1 for OCS, 0.5-1.5 Tg (S) yr–1 for CS2,

6-9 Tg (S) yr–1 for H2S, and 67-100 Tg (S) yr–1 for SO2

(see recent review in Watts, 2000; Penner et al., 2001).

The dominant natural source of sulfur is DMS, of which

about 80-95% is emitted from the oceans. The source for

SO2 consists mostly of anthropogenic sources from fossil

fuel use (~70 Tg (S) yr–1), with volcanic outgassing

(including small eruptions) and biomass burning contributing 8-20 Tg (S) yr–1 (Graf et al., 1997; Andres and

Kasgnoc, 1998) and 2.5 Tg (S) yr–1 (Hao et al., 1990),

respectively. Note that portions of the latter sources are

deposited directly into the free troposphere and can reach

the stratosphere more efficiently. The amount of sulfur

emitted by aviation in the upper troposphere and lower

stratosphere was estimated to be 0.06 Tg (S) yr–1 in 1992

(see Fahey and Schumann et al., 1999). The Mt. Pinatubo

eruption is estimated to have injected 20 Tg (S) directly

into the lower stratosphere.

2.6.2 Chemistry of the VSL Sulfur Species

Here we discuss briefly the chemistry of each of

the VSL sulfur species and some of the current uncertainties. Oxidation of H2S in the troposphere is well understood. It reacts with OH and produces mainly SO2.

Oxidation of CS2 occurs by reaction with OH and possibly by a mechanism involving direct photoabsorption

by CS2, both mechanisms occurring at comparable rates

(Jones et al., 1983). Both routes form OCS and SO2 as

the main stable products. DMS reacts in the gas phase

with OH during the day and with NO3 at night. Figure

2-12 shows the DMS degradation mechanism. SO2 is

the major oxidation product, with dimethyl sulfoxide

(DMSO, CH3S(O)CH3) and methanesulfonic acid (MSA,

VERY SHORT-LIVED SUBSTANCES

2.40

SO

H SO / Sulfate

H SO / Sulfate

OCS CS H S

SO

OH

DMS DMSO MSA

DMS

OH OH NO

OH

Stratosphere

Troposphere

.

OH, O

OH, h

h ν, OH

(O , H O ) aq

OH

 (O )aq

EMISSIONS

DRY AND/OR WET DEPOSITION

TRANSPORT

SEDIMENTATION

OCS

2

2 2

BrO (O )aq

3

3

2

2

2 4

3

2 4

2

ν

3

NO3

3

Figure 2-11. Schematic diagram

of the tropospheric and stratospheric sulfur cycles. The contributions from oxidation of DMS in

the stratosphere have not been

quantified. Those processes are

indicated by dashed arrows.

Production of SO2 from OCS in the

troposphere and direct emission of

H2SO4 are not shown because

their effects on the budget in the

troposphere are small compared

with the other processes.
CH3SO3H) as minor products. The specific yields of DMS

oxidation products, which depend on temperature and

other environmental conditions, are still under debate

(Williams et al., 2001; de Bruyn et al., 2002). Recent laboratory measurements have shown that DMS can also

react with ozone in solution with a rate that is a factor of

about 106 faster than the corresponding homogeneous

gas-phase reaction (Gershenzon et al., 2001). Such a fast

reaction may, under certain conditions, compensate for

the small DMS solubility so that the liquid-phase reactions may be significant in clouds.

The basic chemistry of SO2 oxidation to form

H2SO4 is well understood. It occurs mainly by liquidphase oxidation in clouds or aerosols, involving hydrogen

peroxide (H2O2), O3, or HOX (X = Cl or Br) as oxidizing

agents. The other loss processes for SO2 in the troposphere are dry and wet deposition, and gas-phase oxidation by OH. The gas-phase oxidation represents only

about 10-20% of SO2 loss (e.g., Koch et al., 1999; Barth

et al., 2000; Chin et al., 2000). The combined residence

time of SO2 against all removal processes in the troposphere is estimated to be 1 to 2 days, with a large fraction

of SO2 being removed via deposition and liquid-phase

oxidation. In the stratosphere, SO2 is transformed to

H2SO4 by gas-phase oxidation by OH with a lifetime of

approximately 35 days (Bluth et al., 1992). The sulfate

vapor will either form new particles by homogeneous

nucleation of sulfuric acid and water vapor or condense

on existing aerosol particles.

It is important to keep track of formation of new

particles because smaller particles have longer residence

times in the stratosphere and will lead to a larger mass

loading. There are some indications that the upper troposphere can be a source of new particles, which can then

be transported to the stratosphere where they grow in size

(Brock et al., 1995; Hamill et al., 1997; de Reus et al.,

1999). Transport of small sulfate aerosol particles to the

stratosphere will result in a larger increase in surface area

VERY SHORT-LIVED SUBSTANCES

2.41

Products

CH3S(O)CH3 CH3S(O)OH

O2 OH

DMS(OH)

O2

Multiphase 

Processes

Products

OH OH

H2SO4

OH, O2 SO2 SO3

H2O

∆

CH3SO3H

CH3SCH3

O2

∆

CH3SO3O

OH / NO3 O2

∆

CH3SO2O

CH3SCH2 CH3S CH3SO CH3SO2 CH3SO3

NO NO

NO

O2 / NO O3 / NO2 O3 / NO2

 HO2

RH / HO2

CH3SOO

O3

O3 / NO2

 HO2

∆

O2

∆

O3 (aq)

SO4

=

Figure 2-12. Degradation mechanisms for dimethylsulfide (DMS, CH3SCH3).
than transport of the same mass of sulfur in gaseous form,

which will preferentially condense on existing sulfate particles. The mechanisms and conditions necessary for new

particle formation are still not well understood (de Reus

et al., 2000).

2.6.3 Sources for the Stratospheric Sulfate

Layer

2.6.3.1 DIRECT INPUT FROM EXPLOSIVE VOLCANIC

ERUPTIONS

Direct injections of volcanic sulfur into the stratosphere are sporadic and unpredictable. Explosive volcanic eruptions reach the stratosphere in general at least

once every 2 years (Simkin, 1993). Typically, sulfur

gases, mainly SO2 and H2S, contribute 2 to 35% (in

volume) to the volcanic gas emissions. OCS and CS2

contribute only a small fraction (10–4 to 10–2% in

volume).

Most models simulate the aerosol evolution after a

volcanic eruption assuming that all the sulfur is injected

in the form of SO2. This will not lead to a large error for

the sulfur species because H2S oxidizes to SO2 within 2

days in the troposphere and within 10 days in the stratosphere. The Mt. Pinatubo eruption is estimated to have

deposited 20 Tg (S) in the stratosphere. After the Mt.

Pinatubo eruption, 98% of the observed stratospheric

aerosol was volatile (Deshler et al., 1992; Sheridan et al.,

1992), indicating that homogeneous nucleation is the most

important process for stratospheric aerosol formation in

the disturbed atmosphere. Within the volcanic layers,

Deshler et al. (1992) found an increase of 1-2 orders of

magnitude in the concentration of condensation nuclei.

Model calculations that successfully reproduced Mt.

Pinatubo sulfate aerosol (Timmreck and Graf, 2000;

Weisenstein et al., 1997) were based on homogeneous

nucleation.

Observations (Russell et al., 1996) and model

studies (Zhao et al., 1995; Timmreck and Graf, 2000)

show that it takes about 3 months to build up the sulfate

peak. The volcanically enhanced stratospheric aerosol

layer can be observed for several years after major eruptions. The aerosol optical depth at 550 nm first peaked to

>0.2 three months after Mt. Pinatubo, and then reduced

within 6 years to the background value of 0.003. This represents an e-folding time of approximately 17 months

(McCormick and Veiga, 1992; Ansmann et al., 1997). The

relatively faster removal of large compared with small

particles from the stratosphere leads to a more rapid

decrease in aerosol mass than in aerosol surface (Ansmann

et al., 1997; Godin and Poole et al., 1999).

The presence of the volcanic aerosol would change

the transport circulation in the lower stratosphere (see,

e.g., Kinne et al., 1992; Schoeberl et al., 1993). Most of

the model studies do not take this into account. Questions

have also been raised whether a sufficiently large volcanic

injection could also affect the HOx cycle in the stratosphere. From simple consideration of stoichiometry, it

can be assumed that one molecule of H2O is produced per

H2S oxidized. For SO2, it is assumed that no net gain or

loss of HOx results from SO2 oxidation (Read et al., 1993;

McKeen et al., 1984). Finally, approximately three molecules of H2O are needed to oxidize SO2 and condense

sulfuric acid to sulfate aerosols (Bekki, 1995). Twodimensional model simulations (Bekki et al., 1996)

showed that a hypothetical injection of 6000 Tg SO2 could

dehydrate the stratosphere. Given that most large eruptions inject about several tens of teragrams of sulfur, this

effect is probably not important under most conditions.

2.6.3.2 VSL PRECURSORS EMITTED IN THE

TROPOSPHERE

It has been long recognized that oxidation of OCS

in the stratosphere is an important source of stratospheric

background sulfate aerosol (Crutzen, 1976). Carbonyl

sulfide is emitted at the Earth’s surface by natural and

anthropogenic sources, and is also formed by the oxidation of CS2 (Chin and Davis, 1993, and the references

therein; Watts, 2000). Chin and Davis (1995) used available OCS atmospheric observations and improved laboratory reaction-rate data to show that oxidation of OCS

produces 0.03 Tg (S) yr–1 in the stratosphere. This corresponds to 5% of the OCS emission from the surface and

is only 20-50% of the recently revised estimates of the

sulfur flux (0.06-0.15 Tg (S) yr–1) needed to maintain the

background stratospheric sulfate layer (Chin and Davis,

1995). Later model studies seem to support these conclusions and attribute the other major source to SO2 supplied

from the troposphere. For example, Weisenstein et al.

(1997) found in their 2-D model, which includes a microphysical treatment of the particles, that OCS oxidation

could account for 0.049 Tg (S) yr–1 of sulfate production

in the stratosphere, and that convective transport of SO2

in the tropical troposphere could provide 50% of the background sulfate aerosol. Kjellström (1998) estimated that

the sulfate production from OCS was 3 to 12 times smaller

than the stratospheric background source of sulfate

aerosols, and that the transport of SO2 from the troposphere was the major supplier for the background sulfate

aerosol. A very recent global 3-D model study (Pitari et

al., 2002) estimated that OCS photochemical oxidation

could account for about 43% of the total background stratVERY SHORT-LIVED SUBSTANCES

2.42
ospheric sulfate aerosol, whereas the upward transport of

tropospheric sulfate and SO2 contribute 30% and 27%,

respectively. There are important uncertainties in such

model results because of uncertainties in the model

description of how transport processes control the amount

of VSL substances such as SO2 and sulfate delivered to

the stratosphere (see Section 2.2).

Of particular interest is how changes in anthropogenic SO2 emissions would affect the sulfate loading in

the stratosphere. Chin et al. (2001) and Pitari et al. (2002)

performed simulations to compute the sulfate loading

from anthropogenic SO2 emission. Using the most

recent estimate of emission of 69 Tg (S) yr–1 from

the Intergovernmental Panel on Climate Change

Special Report on Emissions Scenarios (IPCC-SRES)

(Nakic ́enovic ́ et al., 2000), they calculated burdens of 0.7-

0.8 Tg (S) sulfate in the troposphere and 0.07-0.09 Tg (S)

sulfate in the stratosphere. A residence time for sulfate in

the stratosphere of 1-2 years would imply that about 0.05-

0.15% of the SO2 surface emission enters the stratosphere.

No study has been done for SO2 emissions from volcanic

degassing and biomass burning. It is important to

remember that because of different emission locations,

the effect will not necessarily scale with the emission

strengths. Kjellström et al. (1999) estimated that aircraft

emissions contribute less than 1% of the total sulfate mass

in regions of high air traffic.

2.6.4 How Can We Predict Future Changes

in the Stratospheric Sulfate Layer?

If anthropogenic SO2 emissions can account for

more than half of the stratospheric background sulfate

aerosol, future changes in anthropogenic SO2 emission

could have an important impact on stratospheric sulfate

aerosol loading. One can ask whether the past secular

trend in anthropogenic SO2 emission has caused an

increase in the loading of stratospheric aerosol. This question cannot be easily answered from the limited existing

set of observations. Analysis of time series related to the

stratospheric aerosol layer does not show any evidence of

a trend in the nonvolcanic aerosol loading between the

late 1970s, the late 1980s, and the late 1990s to the present

(see Chapter 4). The average global anthropogenic SO2

emission strength has probably been fairly constant

between 1980 and 1990, the increases in some regions

being compensated by decreases in other regions (Boucher

and Pham, 2002). However, changes in the geographical

distribution of anthropogenic SO2 emissions may also

modify the fraction of surface SO2 emissions that contribute to the stratospheric sulfate aerosol layer. Finally,

since the contribution of tropospheric volcanic emissions

to the total atmospheric sulfate burden is nearly equal to

that from anthropogenic sources (Graf et al., 1997), the

climatological stratospheric aerosol levels are also determined by outgassing associated with volcanic activity that

is highly variable in space and time and unpredictable.

Future changes in the stratospheric sulfate layer

have been estimated by performing model simulations

using the IPCC projected emission scenarios for years

2030 and 2100 (Nakic ́enovic ́ et al., 2000). Table 2-13

summarizes the model results of Chin et al. (2001). The

first simulation, SC1, is taken from the IPCC SRES (scenario A2) to represent the anthropogenic emissions in year

2000. Simulations 2 and 3 (SC2 and SC3) are based on

the SRES preliminary emission scenario A2p, which projects a regionally inhomogeneous economic development

for years 2030 and 2100, respectively. Simulation 4 (SC4)

is from another SRES preliminary emission scenario

(B1p) for year 2100, which assumes a convergent world

with a rapid introduction of new and more efficient technologies. It can be deduced from the values in Table 2-13

that every 1 Tg (S) yr–1 increase in anthropogenic SO2

VERY SHORT-LIVED SUBSTANCES

2.43

Table 2-13. Estimated sulfur emission rates in years 2000, 2030, and 2100 and the model-calculated

corresponding tropospheric and stratospheric sulfate burdens. The cited values for the simulated burdens are from Chin et al. (2001).

Emission Rates and Stratospheric SC1 SC2 SC3 SC4

Sulfate Burdens 2000 2030 2100 2100

Anthropogenic emission, SO2 (T (S) yr-1) 69.0 111.9 60.4 28.7

Oceanic emission, DMS (Tg (S) yr–1) 16.1 16.1 16.1 16.1

Volcanic emission, SO2 (Tg (S) yr-1) 4.8 4.8 4.8 4.8

Tropospheric burden (Tg (S)) 0.656 1.124 0.647 0.372

Stratospheric burden (Tg (S)) 0.069 0.143 0.068 0.044

Volcanic emissions are for the continuously erupting volcanoes only and are from Andres and Kasgnoc (1998). Anthropogenic emissions are from four

scenarios in the IPCC SRES (Nakic ́enovic ́ et al., 2000).
emission could roughly result in 9 Gg (S) of sulfate burden

increase in the troposphere and 1 Gg (S) in the stratosphere. Using the emission rates of SC2, Pitari et al.

(2002) have estimated that the stratospheric non-OCS sulfate burden will be 0.14 Tg (S) in 2030, a value similar to

that in Table 2-13.

Using a coarse-grid model with simple chemistry,

Pitari et al. (2002) showed the increase in surface emission of SO2 between year 2000 (SC1) and 2030 (SC2)

will result in an increase in aerosol surface area of up to

0.5 mm2 cm–3 in the Northern Hemisphere, at altitudes

between 100 and 200 hPa. The increase in the stratospheric surface area would cause a slowdown of the

stratospheric ozone recovery rate expected from the reduction in the CFCs.

Overall it can be concluded that representation of

the processes controlling the SSA budget in models is very

much simplified and the calculations of impacts of future

changes in nonvolcanic sulfur emissions are therefore very

uncertain. More studies are needed to understand the relative contributions of surface emission of VSL sulfur substances to the present and future stratospheric sulfate

aerosol burden. In addition to the general problems of

transport of VSL species in the troposphere (discussed in

Section 2.2), areas of uncertainty are the role of aerosol

nucleation in the upper troposphere; the mechanism of

gas-phase DMS oxidation; the heterogeneous reactions of

VSL sulfur substances on aerosol surfaces; the induced

changes in the radiation balance and circulation in the

lower stratosphere; and the magnitude of future emissions.

The attempts by several modeling groups (e.g., Pitari and

Mancini, 2002; Timmreck, 2001; Pitari et al., 2002) represent promising starts. However, such models must be

extended to include a more complete treatment of the

processes in the troposphere and be thoroughly evaluated

with measurements (see, e.g., Lohmann et al., 2001).

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Zuo, Y.G., and Y.W. Deng, Evidence for the production of

hydrogen peroxide in rainwater by lightning during

thunderstorms, Geochim. Cosmochim. Acta, 63,

3451-3455, 1999.

VERY SHORT-LIVED SUBSTANCES

2.57
CHAPTER 3

Polar Stratospheric Ozone: Past and Future

Contributors:

D. Allen

L.M. Avallone

S.R. Beagley

G.E. Bodeker

C. Brühl

J. Christy

B. Connor

M. Dameris

A.R. Douglass

S. Eckermann

M. Gelman

F. Goutail

P. Hamill

Y. Koshelkov

K. Labitzke

U. Langematz

R. Lin

E. Manzini

T. Nagashima

E.R. Nash

J. Nash

S.J. Oltmans

D. Parker

K. Pfeilsticker

G. Pitari

W.J. Randel

E. Rozanov

M.L. Santee

C. Schnadt

M.D. Schwarzkopf

T.G. Shepherd

M. Shitamichi

P. von der Gathen

D.W. Waugh

P.O. Wennberg

Lead Authors:

P.A. Newman

J.A. Pyle

Coauthors:

J. Austin

G.O. Braathen

P.O. Canziani

K.S. Carslaw

P.M. de F. Forster

S. Godin-Beekmann

B.M. Knudsen

K. Kreher

H. Nakane

S. Pawson

V. Ramaswamy

M. Rex

R.J. Salawitch

D.T. Shindell

A. Tabazadeh

D.W. Toohey

CHAPTER 3

POLAR STRATOSPHERIC OZONE: PAST AND FUTURE

Contents

SCIENTIFIC SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.1

3.0 INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.5

3.1 TRENDS OF OZONE AND TEMPERATURE IN THE POLAR STRATOSPHERE . . . . . . . . . . . . . . . . . . . . 3.6

3.1.1 Polar Ozone Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.6

3.1.2 Polar Temperature Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.10

3.2 BASIC POLAR STRATOSPHERIC PROCESSES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.13

3.2.1 Transport and Dynamics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.13

3.2.1.1 The Polar Vortex: Mean Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.13

3.2.1.2 Polar Vortex: Causes of Interannual Variability and Its Implications . . . . . . . . . . . . . . . . . 3.13

3.2.1.3 Polar Transport and Mixing . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.19

3.2.2 Polar Stratospheric Clouds . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.21

3.2.2.1 Observations of PSC Physical Properties and Their Interpretation . . . . . . . . . . . . . . . . . . . 3.21

3.2.2.2 Particle Composition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.24

3.2.2.3 Denitrification . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.25

3.2.2.4 Dehydration . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.28

3.2.3 Polar Ozone Chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.28

3.2.3.1 Chlorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.29

3.2.3.2 Bromine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.33

3.2.4 The Polar Summer Lower Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.37

3.2.4.1 Summertime NOx Chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.38

3.2.4.2 Summertime HOx Chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.39

3.2.4.3 Summertime Cly Chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.40

3.3 QUANTIFICATION OF POLAR OZONE LOSS: OBSERVATIONS AND MODELS . . . . . . . . . . . . . . . . . 3.41

3.3.1 Approaches to Quantify Chemically Induced Ozone Loss in the Arctic . . . . . . . . . . . . . . . . . . . . . . 3.42

3.3.1.1 Approaches That Use Explicit Transport Calculations . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.42

3.3.1.2 Approaches That Use the Relation of Ozone to an Inert Tracer . . . . . . . . . . . . . . . . . . . . . . 3.43

3.3.2 Arctic Ozone Loss During the Last Decade . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.45

3.3.2.1 Ozone Loss Rates Near the Maximum of the Ozone Concentration . . . . . . . . . . . . . . . . . . 3.45

3.3.2.2 Vertical Profiles of Ozone Loss . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.45

3.3.2.3 Effect of Ozone Loss on the Total Ozone Column . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.48

3.3.2.4 Chemical Ozone Loss in the Arctic Winter 1999/2000 . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.50

3.3.3 Consistency Between the Different Observational Techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.50

3.3.4 The Effect of Denitrification on Ozone Loss . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.51

3.3.5 Model Studies of Arctic Ozone Loss . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.54

3.3.5.1 Chemical Transport Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.54

3.3.5.2 Specific Model Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.55

3.3.6 Quantifying Antarctic Ozone Loss . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.57

3.3.7 Model Studies of Antarctic Ozone Loss . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.57

3.3.8 Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.58

3.4 CAUSES OF POLAR STRATOSPHERIC TEMPERATURE TRENDS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.58

3.4.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.58

3.4.2 Modeling Techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.58
3.4.2.1 Fixed Dynamical Heating (FHD) Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.59

3.4.2.2 General Circulation Models (GCMs) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.59

3.4.3 Causes of Polar Stratospheric Temperature Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.59

3.4.3.1 Polar Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.59

3.4.3.2 Well-Mixed Greenhouse Gases (WMGHGs) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.63

3.4.3.3 Combined Changes in WMGHGs and Stratospheric Ozone . . . . . . . . . . . . . . . . . . . . . . . . 3.63

3.4.3.4 Stratospheric Water Vapor . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.64

3.4.3.5 Solar Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.64

3.4.3.6 Volcanoes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.64

3.4.4 Timing of the Springtime Cooling Trend . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.65

3.4.5 Summary and Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.65

3.5 CHEMICAL-CLIMATE MODELING OF THE PAST AND FUTURE POLAR STRATOSPHERE . . . . . . . 3.66

3.5.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.66

3.5.2 The Uncertainties in Chemistry-Climate Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.68

3.5.2.1 Temperature Biases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.68

3.5.2.2 The Simulation of Polar Stratospheric Clouds . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.69

3.5.2.3 The Position of the Upper Boundary of the GCM . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.72

3.5.2.4 Predictions of Planetary Waves . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.73

3.5.2.5 Stratospheric Water Vapor Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.75

3.5.3 Model Assessments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.75

3.5.3.1 The 1960-2000 Time Frame: Ozone Depletion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.77

3.5.3.2 The 2000-2020 Time Frame: Start of Ozone Recovery . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.77

3.5.3.3 The 2020-2060 Time Frame: Complete Ozone Recovery . . . . . . . . . . . . . . . . . . . . . . . . . . 3.79

3.5.3.4 Attribution of Model Ozone Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.80

3.5.4 Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.80

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.81

APPENDIX 3A: SATELLITE MEASUREMENTS IN THE ANTARCTIC AND ARCTIC . . . . . . . . . . . . . . . . 3.104
SCIENTIFIC SUMMARY*

• Springtime Antarctic ozone depletion remains very large. Severe depletion, leading to minimum values around

100 (Dobson units (DU), has been seen every year since the early 1990s. The main processes are well understood.

• Some estimates of the severity of the Antarctic ozone hole, e.g., the area enclosed by the 220-DU contour,

show an increase at some times in recent years, and therefore it is not yet possible to say that the ozone hole

has reached its maximum. Much of the change appears to be associated with processes at the edge of the polar

vortex and is consistent with meteorological variability and the almost constant halogen loading. The observations

do not show evidence of ozone recovery; this is expected due to the slow decrease of stratospheric chlorine and

bromine to pre-ozone-hole levels (see Chapter 1).

• Meteorological observations show that the Antarctic polar vortex is persisting later than was observed during

the 1970s. Over the last decade, the vortex has broken up in the early-December period in contrast to a breakup in

late-November during the 1970 to 1980 period.

• There was very large local ozone depletion in the Arctic vortex in 1999/2000, reaching 70% by early April in

a narrow region around 20 km. Integrated column losses were greater than 80 DU. The winter of 1999/2000 was

characterized by persistent low temperatures and a strong vortex. In contrast, in the warmer more disturbed polar

vortex of 1998/1999, the estimated loss was very small. These observations are consistent with our expectation that

Arctic ozone losses are largest in cold stratospheric winters.

• The Arctic winter/spring ozone column continues to be variable, reflecting the variable meteorology of the

Northern Hemisphere (NH) stratosphere. Lower column ozone was present during the cold winter of 1999/2000

than in the warmer, more disturbed winters of 1998/1999 and 2000/2001, reflecting the variability in the dynamical

and chemical processes that control the ozone layer in the NH stratosphere. It is not possible to isolate the importance of these factors just from observations; model studies are needed to do this. The Arctic Oscillation can be

used as an index to describe variability, but not causality.

• The magnitude of chemical loss of ozone for all Arctic winters during the last decade has now been studied

with a variety of observationally based approaches. There is generally good agreement between different

analyses for quantifying losses: for the 1999/2000 winter, agreement was better than 20% in the Arctic stratosphere

around 20 km.

• Satellite and radiosonde observations show that the springtime Arctic and Antarctic lower stratospheres

have cooled. However, because of large variability in the Arctic spring, the magnitude of the trend is uncertain

there. During the 1979 to 2000 period the linear temperature trend exceeds –1.5 K/decade at 70°N and 70°S.

• Modeling studies now demonstrate that the stratospheric ozone depletion has exerted an important influence

on the springtime cooling of the Arctic lower stratosphere over the 1980 to 2000 period, but the degree of

attribution is hindered by the large dynamical variability in this region. In Antarctica modeling studies reaffirm that ozone loss is the major cause of the springtime cooling and the increased persistence of the Antarctic

polar vortex. Well-mixed greenhouse gases and stratospheric water vapor increases also contribute to the annually

averaged cooling.

• A laboratory study of the rate of formation of ClOOCl (the chlorine monoxide dimer), the initial step for the

most important polar ozone loss cycle, suggests it is up to 25% faster (depending on temperature) than previous estimates based on data extrapolations to low polar temperatures. This leads to faster calculated rates of

chemical ozone loss.

POLAR OZONE

3.1

* Editor’s note added in press: The preparation and review of this Assessment report were completed by August 2002. The observations of the unusual

Antarctic ozone hole of 2002 occurred thereafter and hence are not included here. Such subsequent observations will, of course, be included in future

reports and assessments.
• In situ observations of the rate of decay of chlorine monoxide (ClO) at sunset are consistent with the new

laboratory rate for ClO + ClO + M and recommended cross sections for ClOOCl. This suggests that errors in

model representation of the partitioning of ClO and its dimer cannot account for discrepancies between modeled

and measured Arctic ozone losses observed in January.

• Modeling studies of the latitudinal, seasonal, and diurnal variations in bromine monoxide (BrO) column

abundances agree well with observations from a number of ground sites, indicating that the processes that

govern bromine partitioning in the polar regions are reasonably well understood.

• Observations of BrO in the winter Arctic vortex by in situ and remote detection techniques are in broad

agreement and consistent with a total bromine budget of ~20 ± 4 parts per trillion. This result now allows for

more accurate assessment of the contribution of bromine to polar ozone loss. At present, the fractional contribution

of bromine to total ozone loss ranges between 30% and 60%, depending on temperature and abundances of ClO.

Considering the observed leveling off of abundances of sources of chlorine (reported in Chapter 1), the role of

bromine in polar ozone loss will continue to increase relative to that of chlorine until the current upward trends of

the bromine source gases reverse.

• New laboratory and field studies have led to refinements in the recommendations for the rate constants of

several key reactions that couple the photochemistry of odd hydrogen (HOx) and nitrogen oxides (NOx)

(species that are largely controlled by natural processes) and to the discovery of a new process (near-infrared

photolysis of peroxynitric acid (HNO4)). Together with new observations of HOx, NOx, and ozone in late spring

and summer, these studies have demonstrated that our understanding of the photochemistry of HOx and NOx in the

lower summertime stratosphere is fundamentally sound.

• Removal of nitrogen compounds (denitrification) has been observed to occur in the Arctic lower stratosphere

in several cold winters. Denitrification of up to 70% of the total reactive nitrogen was observed at some levels of

the lower stratosphere in winter 1999/2000. Observations and modeling results show that denitrification in the

1999/2000 Arctic lower stratosphere increased ozone loss by as much as 30% at 20 km in spring.

• Our understanding of what causes denitrification has improved considerably by the discovery of large nitricacid-containing particles in the Arctic polar lower stratosphere during the winter of 1999/2000. Sedimentation

of these particles can account for observed Arctic denitrification, although the mechanism of formation of these

sedimenting particles remains uncertain. Sedimentation of ice containing dissolved nitric acid, which has been the

preferred mechanism in stratospheric models, was not the dominant mechanism in the Arctic in 1999/2000. These

observations show that denitrification can occur at higher temperatures than previously thought.

• Synoptic and mesoscale motions (baroclinic and gravity waves) can lead directly to, and enhance, polar

stratospheric cloud (PSC) formation in both hemispheres. For the first time, operational meteorological analyses

have been demonstrated to contain credible information about the gravity-wave field in high latitudes.

• Model calculations suggest that the magnitude and vertical extent of denitrification could increase considerably in a future colder Arctic stratosphere, leading to increased ozone loss over a broader altitude range in

the lower stratosphere. The denitrification mechanism is not well represented in current global models, which is

one of the limiting factors in the ability of the models to reproduce the large ozone losses observed in cold Arctic

winters and to reliably predict future ozone losses in the Arctic.

• The chemical composition of liquid and solid PSC particles has been measured directly for the first time.

Measured compositions are in agreement with model calculations for liquid particles and nitric acid trihydrate,

which have been used in stratospheric models for many years. These measurements give confidence in the microphysical models that are central to simulations of polar ozone loss.

POLAR OZONE

3.2
• Significant chemical loss of ozone (~0.5 parts per million by volume) in the lower stratosphere during January

has been observed in several cold Arctic winters. The observations indicate that the loss occurred exclusively

during periods when the air masses are exposed to sunlight. These January ozone losses cannot be fully explained

with our current understanding of the photochemistry. For some cold Arctic winters the ozone loss during January

contributes about 25% to the overall loss of ozone over the winter.

• Coupled chemistry-climate simulations broadly reproduce past trends in total ozone over the Antarctic. The

coupled chemistry-climate models have studied the future evolution of minimum Antarctic ozone. These models

suggest that the minimum column ozone may have already occurred or should occur within the next decade, and

that recovery to 1980 levels may be expected in the 2045 to 2055 period. The model response is driven mainly by

the changes in stratospheric halogen loading, with a small delay in ozone recovery due to cooling of the lower stratosphere.

• The area of the Antarctic ozone hole depends on processes near the edge of the polar vortex. Area estimates

amongst the coupled chemistry-climate models are uncertain because vortex-edge temperatures are near the

threshold for PSC formation and these temperatures are difficult to predict.

• Coupled chemistry-climate models now capture the typical interannual variability of Arctic ozone levels.

However, temperatures are often near the threshold for PSC formation and, hence, the initiation of perturbed chemistry. This places severe constraints on model predictions of past and future ozone behavior. A number of coupled

chemistry-climate models run for this Assessment suggest that minimum Arctic ozone would occur within the next

two decades, depending on the meteorology. Very low Arctic ozone columns, similar to those seen in the Antarctic,

are not predicted by these models (in contrast to earlier simpler calculations considered in the previous Assessment).

Such low levels in the Arctic would require weak dynamical forcing that is unprecedented in NH observations.

POLAR OZONE

3.3

POLAR OZONE

3.5

3.0 INTRODUCTION*

This chapter provides an update on our understanding of recent changes in polar ozone and of the polar

vortex, and considers possible future developments. It

builds on earlier Assessments, concentrating mainly, but

not exclusively, on work reported since the previous

United Nations Environment Programme/World Meteorological Organization (UNEP/WMO) assessment report

(WMO, 1999).

The previous Assessment reported that the

Antarctic ozone hole continued unabated, with essentially

near-complete destruction of ozone in late winter/spring

in the lower stratosphere, and that the factors controlling

the depletion (meteorological preconditioning, halogen

activation, ozone depletion in sunlight) were well understood. In the Arctic, substantial ozone losses were

reported in several winters during the 1990s, depending

on the meteorological conditions. The Assessment highlighted the vulnerability of the Arctic to large ozone losses

in a cold winter while chlorine abundances remain high

during the next decade or so. Less chemical loss was to

be expected in the Arctic in winters with a warm, disturbed

vortex. Difficulties with the precise quantification of

Arctic ozone loss were indicated. The previous

Assessment highlighted specific uncertainty issues surrounding the understanding of the different types of polar

stratospheric clouds (PSCs), and the process of denitrification, which can limit our ability to model present and

future polar ozone loss.

The coupling between atmospheric chemistry and

climate has been recognized increasingly in recent assessments. In WMO (1999) a late-winter/springtime cooling

in the Arctic polar lower stratospheric temperatures of ~3

to 4 K/decade was noted (although with the large dynamical variability in that region the statistical significance of

the trend was not high), and the role of ozone, water vapor,

and the well-mixed greenhouse gases was explored.

Three-dimensional (3-D) coupled chemistry-climate

models were used for the first time to look at the possible

recovery of the ozone layer; these models all indicated a

delay in recovery beyond the time of the peak in stratospheric halogen abundance.

Since the previous Assessment there has been considerable progress in basic research that we report below.

Satellite datasets on ozone and temperature have been further extended. In addition, scientific impetus has been

provided by several major field campaigns to study the

Arctic stratosphere. Results are reported here from the

National Aeronautics and Space Administration (NASA)

Photochemistry of Ozone Loss in the Arctic Region in

Summer (POLARIS) campaign, aimed at understanding

the summer polar stratosphere; the European Union (EU)

Third European Stratospheric Experiment on Ozone

(THESEO), a polar and middle-latitude campaign; and

the joint NASA/EU Stratospheric Aerosol and Gas

Experiment (SAGE) III Ozone Loss and Validation

Experiment (SOLVE)-THESEO 2000. These campaigns

produced new data to address some of the uncertainties

remaining after the previous Assessment.

Section 3.1 updates polar ozone measurements in

both Antarctica and the Arctic, concentrating on the winter

and spring seasons when the largest ozone depletion is

observed. The total ozone columns are considered, and

updated information on various possible indicators of

ozone recovery, suggested in the previous Assessment, is

presented briefly. The updated polar temperature trends

are also presented here.

Section 3.2 reviews our understanding of the relevant physical and chemical processes controlling the polar

vortex and its composition. The Arctic field campaigns

have provided new data on a disturbed winter with considerable exchange between polar and middle latitudes

(1998/1999) and on the cold polar winter of 1999/2000,

which led to large local ozone depletion. Many new complementary constituent measurements provide an important constraint on chemical loss processes. Important new

measurements of particles were also made in the winter

polar stratosphere, leading to advances in our understanding of particle composition and denitrification.

Improved understanding of the dynamics in and around

the polar vortex has also been developed.

Section 3.3 looks in detail at our quantitative understanding of polar ozone loss. In earlier assessments it was

recorded that models often fail to quantify correctly the

observed ozone loss. A variety of methods to derive ozone

loss from measurements are reviewed in this section.

Estimated losses in recent Arctic and Antarctic winters

are considered and compared with each other and with

model estimates.

In the previous Assessment (WMO, 1999), it was

recognized that the future development of the ozone layer

does not depend just on changes in stratospheric halogen

loading but also, very importantly, on a number of other

factors connecting chemistry and climate. These factors

are discussed in Section 3.4. Temperature changes are

* Editor’s note added in press: The preparation and review of this Assessment report were completed by August 2002. The observations of the unusual

Antarctic ozone hole of 2002 occurred thereafter and hence are not included here. Such subsequent observations will, of course, be included in future

reports and Assessments.
POLAR OZONE

3.6

particularly important, since polar heterogeneous chemistry is strongly temperature dependent, and furthermore,

temperature changes are related to the strength of the polar

vortex, descent within the vortex, and mixing with lower

latitudes. The attribution of the trends in polar stratospheric temperatures (presented in Section 3.1.2) is discussed, and the roles of changes in well-mixed greenhouse

gases, ozone, water vapor, and aerosol particles are

reviewed. Future stratospheric temperature changes are

discussed.

Finally, in Section 3.5, possible future states of the

polar stratosphere are explored in sensitivity calculations

using coupled chemistry-climate models. Results from

these models were reported for the first time in an assessment in WMO 1999, and the models are still being developed. An extensive review of the present uncertainties in

chemistry-climate models is presented here, followed by

some examples of sensitivity calculations to consider the

polar stratosphere during the next 50 years.

3.1 TRENDS OF OZONE AND TEMPERATURE

IN THE POLAR STRATOSPHERE

3.1.1 Polar Ozone Trends

Ozone is primarily produced in the midlatitudes

and tropics by photodissociation of oxygen by hard ultraviolet (UV) radiation (below 242 nm) and is transported

toward the poles by the Brewer-Dobson circulation. The

result is an annual cycle in ozone, shown by the climatological values in Figure 3-1. Because of the stronger

Brewer-Dobson circulation in the Northern Hemisphere

(NH), the Arctic is both warmer and has larger column

ozone amounts than the Antarctic. In the NH, there is usually a maximum in the column in late winter/early spring.

At the South Pole, there is less annual variation (larger

annual variations are expected at the vortex edge). In

recent years, the annual cycle has been modified by polar

ozone depletion, most obviously in the Southern

Hemisphere (SH).

Figure 3-1 also shows recent year-round ozone

measurements from the Arctic Ny Ålesund station

(78.9°N, 11.9°E) and the South Pole station, updating

polar observations since the previous Assessment. The

Antarctic observations in the last few years continue to

show the extremely low spring ozone values that have

characterized the ozone hole during the 1990s. The low

Antarctic values begin with the chemical ozone losses

during August and September and end upon the breakup

of the vortex in November or December. During the Arctic

summer, ozone is destroyed photochemically, especially

at these high latitudes during continuous sunlight conditions (Brühl et al., 1998), and the climatological seasonal

minimum is reached in autumn.

In the Arctic, the March-to-April ozone maximum

is occasionally reduced below the climatology in some

years (e.g., in 1997, the cyan triangles in Figure 3-1)

because of severe chemical ozone loss and reductions in

ozone transport (Andersen and Knudsen, 2002). In these

low-ozone years, the column ozone rapidly increases with

the breakup of the vortex (e.g., early April 1997). Of the

most recent winters, 1999/2000 also has somewhat lower

ozone columns than the climatology, as discussed later.

Extremely low Ny Ålesund column ozone values at the

beginning of 1996 (Figure 3-1, dark red triangles) can partially be explained by the early onset of the ozone depletion that year (see Section 3.3.2). The lowest 1996 values

occur in an “ozone mini-hole” event (Weber et al., 2002)

(see Section 3.2.1.2). A complicating factor in the upper

panel in Figure 3-1 is that occasionally Ny Ålesund is

outside the vortex, as is evident from the large variations

on a broad range of spatial and temporal scales.

Extremely high Ny Ålesund column ozone values

in December 1998 (Figure 3-1, gray diamonds) and

February 2001 (red circles) were caused by sudden warmings and the associated ozone transport. The warmings in

those years also resulted in high temperatures, thereby

preventing the formation of polar stratospheric clouds.

The largest ozone depletion occurs in the polar vortices during springtime. Figure 3-2 shows the springtime

ozone values in the Arctic and Antarctic (63° to 90°) since

1970 (updated from Newman et al., 1997). The Arctic

column ozone averages were extremely low during the

mid-1990s, but have been relatively high in four of the

last five winters. As noted in the previous paragraph and

as is apparent in Figure 3-1, these higher ozone values are

associated with stratospheric sudden warmings. The

downward secular ozone trend apparent through 1997 and

its reversal over the last few years can be associated with

a long-term variation of stratospheric warmings.

Figure 3-2 also shows that the Antarctic ozone hole

continued to display the low values over the last 4 years

that were apparent during the early and mid-1990s. The

notably higher value in October 2000 resulted from greater

dynamical activity, as is also apparent in Figure 3-1.

The polar column ozone averages of Figure 3-2 in

the 63° to 90° region generally coincide with the polar

vortices. However, in the NH the vortex is usually smaller

and the 63° to 90° region may contain air outside the

vortex. The absolute minimum in the NH occurred in

1997, when the vortex was cold, very large, and polecentered. However, the column chemical ozone loss in

the vortex was probably larger in 1995, 1996, and 2000

(see Section 3.3.2.3). However, Andersen and Knudsen
POLAR OZONE

3.7

100

150

200

250

300

350

400

450

500

550

600

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec

100

150

200

250

300

350

400

450

500

550

a) Ny-Ålesund(78.9 N,11.9 E) Column ozone (DU)

TOMS 79-83mean

1994 1995

1996 1997

1998 1999

2000 2001

b) SouthPole Column ozone (DU)

Month

1967-1971climatology

1998 1999

2000 2001

1970 1975 1980 1985 1990 1995 2000

200

250

300

350

400

450

500 Total ozone (DU) BUV

Nimbus-7

Meteor-3

ADEOS

EP-TOMS

NOAA-14

NOAA-9

SH October

NH March

Combined changes due to

chemistry and dynamics

1970-82 Avg.

Figure 3-2. Average total ozone poleward of 63° latitude in March in the

Northern Hemisphere (NH) and October

in the Southern Hemisphere (SH). Symbols indicate the satellite data that have

been used in the different years. The horizontal gray lines represent the average

total ozone for the years prior to 1983 for

the NH and SH. The lighter gray shading

shows the combined differences resulting

from chemical losses and dynamical

processes. Updated from Newman et al.

(1997).

Figure 3-1. The column ozone determined from (a) ozonesondes launched from Ny Ålesund and (b)

ozonesonde and Dobson measurements at South Pole, as a function of day of year. Symbols give individual

measurements. The solid line in panel (a) is the daily TOMS mean column ozone for the Ny Ålesund station

from 1979 to 1983, with the shading indicating the range. The solid line in (b) is a climatology for the South

Pole from 1967 to 1971 (with 1s error bars indicated). Courtesy Peter von der Gathen (Alfred Wegener

Institute) and Samuel Oltmans (NOAA-CMDL).
POLAR OZONE

3.8

(2002) have argued that about 75% of the 63° to 90°N

depletion from 1992 to 2000 relative to the 1979-to-1982

average is due to ozone depletion inside the vortex, so the

plot does give a good indication of the Arctic vortex depletion.

Large total column ozone trends have been seen in

both the Arctic and Antarctic polar vortices during the

spring (Figure 3-3). To obtain these trends with better

correlation with the polar vortices, a potential vorticity

coordinate (equivalent latitude) remapping technique was

applied to a trend analysis of homogenized satellite data

from the Total Ozone Mapping Spectrometer (TOMS) and

the Global Ozone Monitoring Experiment (GOME)

(Bodeker et al., 2001; see also Appendix 3A for satellite

data descriptions). In this coordinate the centers of the

vortices are at 90° and the edges at about 60° to 75° (as

shown by the ¥ points in Figure 3-3). After a regression

model including trends and variability (seasonal cycle,

quasi-biennial oscillation (QBO), solar cycle, volcanic

effects, and El Niño-Southern Oscillation (ENSO)) was

applied to the data from 1978 to 1998, statistically significant linear trends were obtained.

The largest Arctic negative trend (1.04 ± 0.39%

yr-1) is observed in March, whereas the largest Antarctic

negative trend (2.51 ± 0.62% yr-1) is observed in October

(updated from Bodeker et al. (2001) to include the additional years 1999 and 2000). In the Arctic vortex the 1978

to 1998 trend was largely due to severe vortex depletions

in the 1990s (Section 3.3.2). The addition of the warmer

winter of 1998/1999, when no significant ozone depletion

in the vortex occurred, reduced the downward trends. In

the Antarctic vortex trends have weakened due to saturation of the ozone losses. Hence, the trends, with 1999 and

2000 data included, have slightly smaller negative trends

then previous estimates in the Antarctic spring column

ozone. In the Arctic, Figure 3-3 shows steep gradients in

the trends across the edge of the vortex. It also reveals

statistically significant negative ozone trends in May,

June, and July just inside the Antarctic vortex, which were

not found in previous trend analyses, but confirms earlier

findings (Roscoe et al., 1997; Lee et al., 2001) (see Section

3.3.7).

Because the ozone loss mainly occurs in the 12- to

20-km layer, the partial column ozone from 12 to 20 km

Month

-90

-60

-30

30

60

90 Equivalent latitude (550K)

-2.8 -2.2 -1.6 -1 -0.7 -0.4 -0.1

Trend in %/year

Jul Aug Sep Oct Nov Dec Jan Feb Mar Apr May Jun

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec

Figure 3-3. Total column ozone trends as a function of equivalent latitude and season using corrected TOMS

and GOME data from November 1978 to December 2000. Gray shaded regions indicate where trends are

statistically insignificant (at 95% confidence). The crosses show the mean position of the vortex (remnant)

edge(s) over the time period when the position could be determined, and the “error bars” on each cross show

the mean width of the vortex boundary (Nash et al., 1996). The equivalent latitude is calculated on the 550-K

isentropic surface (Bodeker et al., 2001).
provides a good representation of the long-term decrease

of the stratospheric ozone in Antarctica and could be used

as an indicator of ozone recovery (see WMO, 1999).

Figure 3-4 gives monthly averaged partial column ozone

in September, October, and November (SON) based on

ozonesonde observations at Syowa (1968 to 2001). The

partial column ozone has decreased considerably from the

early 1970s (~80% in September, ~85% in October, ~80%

in November). The October partial column has not shown

appreciable change since 1992. On the other hand, the

September and November partial columns continued to

show small reductions during the 1990s. The averaged

partial column ozone over the 3 months (September to

November) has also shown decreases during the 1990s,

with relatively smaller interannual variability. Because

Syowa is located near the vortex edge region, these

September and November ozone reductions during the

1990s may be related to cooling near the vortex collar.

Spatially averaged characteristics of the Antarctic

ozone hole from 1979 to 2001 based on the total column

ozone observed by satellites (TOMS series for 1979-2001,

and Television Infrared Observation Satellite (TIROS)

Operational Vertical Sounder (TOVS) for 1995) are shown

in Figure 3-5. These parameters include the maximum area,

the minimum total column ozone, the ozone mass deficiency, and the date of the ozone hole’s disappearance. The

maximum area of the ozone hole increased rapidly during

the 1980s and gradually during the 1990s, with year-to-year

variations, and reached a maximum in 2000. The minimum

total ozone, which usually appears in late September or in

early October, has been approximately 100 Dobson units

(DU) since 1993 (Figure 3-5) after the considerable

decrease during the 1980s and the early 1990s. The ozone

mass deficiency in the ozone hole (O3 MD) is defined as

the ozone mass deficiency from 300 DU in the sunlit area

poleward of 60°S averaged for 105 days (1 September to

15 December). O3 MD varied in concert with the Antarctic

ozone hole area and was at the highest level ever in 2000.

The date of the disappearance of the Antarctic ozone hole

(disappearance of the total ozone values below 220 DU)

has generally been occurring later in the season. As a whole,

observations show that the Antarctic ozone hole has been

slightly larger in the last few years in comparison with the

mid-1990s. These observational results could be explained

by ozone decreases near the vortex edge (e.g., Bodeker et

al., 2001, 2002; Lee et al., 2001). Although the size of the

Antarctic polar vortex has not increased, it has been

stronger, as shown in Figure 3-6. There has been a tendency toward a cooling of the vortex due to ozone depletion, and the polar vortex has been more persistent, with

some interannual variability (the temperature trends and

persistence of the Antarctic polar vortex are discussed in

the next section, Section 3.1.2). These conditions could

result in more extensive polar stratospheric clouds (PSCs)

in the sunlit vortex edge, and larger chemical depletion of

ozone. This expands the area of the hole, and delays its

disappearance into the late-spring period.

POLAR OZONE

3.9

0

50

100

150

200 Ozone (DU)

1965 1970 1975 1980 1985 1990 1995 2000

September

1965 1970 1975 1980 1985 1990 1995 2000

1965 1970 1975 1980 1985 1990 1995 2000

1965 1970 1975 1980 1985 1990 1995 2000

0

50

100

150

200 Ozone (DU)

0

50

100

150

200 Ozone (DU)

0

50

100

150

200 Ozone (DU)

October

November

Average (Sep.-Nov.)

Figure 3-4. Partial column ozone in the 12- to 20-

km layer in September, October, and November, and

the average for September-November, over Syowa

in Antarctica from 1968 to 2001. Update of Figure

4-27 of WMO (1999) by Japan Meteorological

Agency.
The ozone hole has expanded in area somewhat

since the early 1990s. The top panel of Figure 3-5 displays the maximum area of the ozone hole observed in the

1979 to 2001 period. There appears to be a steady increase

of ozone hole size since the early 1990s. However, maximum values are sensitive to the timing of individual

dynamical events in the early-September period (e.g., the

30 million km2 single-day value in 2000). Figure 3-7 displays the average size of the ozone hole (determined by

the area enclosed by the 220-DU total ozone contour) over

this same period, as determined from TOMS observations

(see Appendix 3A for TOMS data description). Again,

larger sizes are observed in the 1998 to 2001 period as

compared with the 1992 to 1997 period. These size estimates are generally consistent with predictions that the

ozone hole would have slight growth over the 1992 to

1995 period and maximum sizes occurring in the period

after 1996 (Schoeberl et al., 1996).

3.1.2 Polar Temperature Trends

Substantial observational temperature data on the

polar stratosphere are available from ~1979, including

radiosonde and satellite measurements, and analyses of

various types (Ramaswamy et al., 2001). Figure 3-8 illustrates the time series of temperatures at 70°N (March) and

70°S (November) from National Centers for Environmental Protection (NCEP) reanalyses and Climate

Prediction Center (CPC) analyses (for descriptions, see

WMO, 1999, Chapter 5). There are large interannual variations manifest in both hemispheres that generally complicate the determination of statistically significant trends.

POLAR OZONE

3.10

0 15 30 45 60 75 90

Equivalent latitude (deg)

0

5 0

100

150

200

250

300

350

400 Zonal mean ozone (DU)

0

100

200

300

400

500

600

700 PV grad x w ind speed (K m3 s- 2 k g-1 deg- 1) 1981-1985 1986-1990 1991-1995 1996-2000

1996-2000

1991-1995

1986-1990

1981-1985

9

10

11

12 Date of disappearance Date of disap earance of the ozone hole (month) of the ozone hole (month)

1980 1985 1990 1995 2000

30

25

20

15

10

5

0 Area (10 Area (106 km2) As compared with the area of As compared with the area of Antarctica Continent Antarctica Continent

1980 1985 1990 1995 2000

2.0

1.5

1.0

0.5

0.0

1980 1985 1990 1995 2000 Mininum Total Ozone Mininum Total Ozone (Dobson units) (Dobson units) 300

250

200

150

100

50

0 Ozone Mass Deficiency Ozone Mass Deficiency (106 tons) 100

80

60

40

20

0 1980 1985 1990 1995 2000

(a)

(b)

(c)

(d)

Figure 3-5. Evolution of the Antarctic ozone hole

(defined by total ozone values below 220 DU) based

on satellite data for the period 1 September to 1

December of each year. Panel (a) shows the maximum area covered by ozone hole values (left axis

in millions km2; right axis gives size relative to

Antarctic area), (b) the minimum total column ozone

measured each year, (c) the maximum ozone mass

deficiency from 300 DU during the period 1

September to 15 December (in Mt), and (d) the date

of disappearance of the ozone hole values (Uchino

et al., 1999; data updated to 2001).

Figure 3-6. October SH average zonal mean total

column ozone (top four thick solid lines plotted

against the left ordinate) and October average vortex

strength (gradient of potential vorticity (PV) multiplied by zonal wind speed; thin solid lines plotted

against the right ordinate) for each of the four

analysis periods (Bodeker et al., 2002).
POLAR OZONE

3.11

The Microwave Sounding Unit channel-4 (MSU4) data and the Stratospheric Sounding Unit (SSU) derived

temperature trends (1979 to 1998) for 70°N and 70°S are

shown in Figure 3-9 (see also Figures 3-39 and 3-40).

These are an update to the 1979 to 1994 trends presented

in Ramaswamy et al. (2001). The 70° latitude is chosen

for comparison, because this is the highest latitude for

which SSU trend data are available. Both MSU-4 (black

line) and SSU-15X (gray line) signals originate from a

range in altitude in the stratosphere and do not correspond

to one particular height. The SSU peak signal corresponds

to a pressure of roughly 50 hPa and shows a statistically

significant (at the 2s level) cooling of nearly 3 K/decade

at 70∞N in March, April, and May (MAM) and 70∞S in

September, Ocotber, and November (SON), and approximately 1.2 K/decade for the annual average temperature

change at both poles. The instrument also shows cooling

significant at the 1s level for most other seasons. The

MSU-4 data (peak signal from approximately 100 hPa)

also shows a significant cooling during the spring in both

hemispheres (–1.8 K/decade at 70°N and –1.1 K/decade

at 70°S) and at both poles. These MSU-4 trends are

roughly half the magnitude of the SSU trend. The

observed satellite trends for this period indicate cooling

in all seasons at both 70°N and 70°S.

It should be noted that the magnitude and statistical significance of the trends in both regions are

dependent on the end-year considered. This is more crucial for the Arctic, especially during winter/spring when

1980 1985 1990 1995 2000 0

5

10

15

20

25

30 Ozone Hole Area (Million k m2

)

Antarctica area

Figure 3-7. Area of the Antarctic ozone hole

(defined by total ozone values below 220 DU) based

on TOMS satellite data. The points show the

average of the daily areas (in millions km2

) between

7 September and 13 October. The vertical lines

show the range of values over this same period.

March 70°N

November 70°S

19651970197519801985199019952000

19651970197519801985199019952000

20

10

0

-10

-20

20

10

0

-10

-20

Temperature (K) Temperature (K) Figure 3-8. Zonal mean temperature deviation from

the long-term mean in the lower stratosphere during

March at 70°N (top) and November at 70°S (bottom).

The red filled circles are 50-hPa NCEP/CPC objective analysis data; the open red circles are from the

50-hPa Freie Universität Berlin analyses; the open

green circles are the MSU channel-4 satellite data

(approximately a layer mean between 12 and 22

km); the filled green circles are the SSU channel15X data (approximately a layer mean between 12

and 28 km); the open blue circles are the UK

Meteorological Office adjusted and gridded

radiosonde (RAOB) data at 50 hPa, and the filled

blue circles are 50-hPa Russian radiosonde data

from high northern latitudes. For each dataset, the

deviation was determined with respect to the

monthly mean of the time series (see Ramaswamy

et al. (2001) for further details on the datasets).

Figure assembled in cooperation with the Stratospheric Processes and Their Role in Climate

(SPARC) Stratospheric Temperature Trends

Assessment project.
POLAR OZONE

3.12

the time series reveals large interannual variations in temperatures (see Figure 3-8; also Labitzke and Van Loon,

1995). The trend sensitivity can be appreciated by comparing the latest (1979 to 1998) MSU trend at 70°N

(March, as shown in Figure 3-9) with the corresponding

MSU trends shown in WMO (1995, Figure 8.11 for the

period 1979 to 1991).

Comparison of the satellite 1979 to 2000 annualmean trends with those obtained for the 1979 to 1994

period (WMO, 1999) shows that there is now a statistically significant cooling at the 95% confidence level in

the mid-to-high southern latitudes (Ramaswamy et al.,

2002a). The northern midlatitudes continue to exhibit a

statistically significant cooling trend (see WMO, 1999)

while the higher latitudes (Arctic region) now have a

cooling trend significant at the 90% confidence level. As

in the satellite data, the 1979 to 2000 sonde trends yield

an annually averaged cooling trend in the northern polar

region. The sonde trend is somewhat smaller than the

satellite trend (see Figure 3-39), although this may be partially due to the time period for the trend analysis being

longer in the sonde data (note that 3 of the last 4 years

have been relatively warmer; see Figure 3-8). The CPC

analysis at 50 hPa also shows a cooling at both poles, consistent with the satellite data and sonde data.

Because ozone is radiatively active in both the

infrared and the ultraviolet, large chemically driven polar

ozone losses can potentially alter the dynamics of the stratosphere by cooling the polar stratosphere (Kiehl et al.,

1988; see also WMO, 1990). Randel and Wu (1999a) have

shown such an Antarctic cooling using radiosonde and

conventional meteorological analyses, and Compagnucci

et al. (2001) using MSU retrievals. This cooling

strengthens the meridional temperature gradient, thereby

strengthening the Antarctic polar vortex. Waugh et al.

(1999) and Zhou et al. (2000) have additionally shown

that the Antarctic vortex has strengthened over the last

two decades, with breakup dates occurring later in the

spring, consistent with the polar cooling forced by the

ozone loss. Figure 3-10 displays these stratospheric polar

40

60

80

100

120 Day No.

Feb

Mar

Apr

1950 1960 1970 1980 1990 2000 2010

300

320

340

360

380

400 Day No.

Nov

Dec

Jan

NH

Reanalysis SH

CPC

1950 1960 1970 1980 1990 2000 2010

-5

-4

-3

-2

-1

0 Temperature trend (K/decade)

70°N

70°S

MSU4

SSU15x

-4

-3

-2

-1

0

2

Jan Feb Mar Apr May Jun Jul AugSepOct Nov Dec

Figure 3-10. The vortex breakup date from 1958 to

2002 on the 500-K surface in the Northern (top) and

Southern (bottom) Hemispheres using the method

described in Nash et al. (1996; see also Waugh et

al., 1999). The thin lines with filled dots are the dates

calculated from the National Centers for Environmental Prediction/National Center for Atmospheric

Research (NCEP/NCAR) reanalysis data. The thick

lines are these same data time-filtered to remove

the year-to-year fluctuations. The crosses are the

breakup dates calculated from the NCEP/Climate

Prediction Center analyses for the period 1979 to

2002.

Figure 3-9. Stratospheric temperature trends for

the period 1979 to 1998 at 70°N (top panel) and

70°S (bottom) for MSU-4 (black line) and SSU-15X

(gray line) for each month of the year. The error bars

represent the 95% confidence limits of the trend estimates. Figure provided by W.J. Randel (NCAR).
POLAR OZONE

3.13

vortex breakup dates in the NH (top) and SH (bottom), as

determined using a wind average along the edge of the

polar vortex (see Nash et al., 1996). The SH breakup date

shows variations of 1 to 2 weeks, with a trend from late

November in the 1980s to mid-December over the last

few years. The NH breakup dates also show considerable

variation, but without a linear trend. As discussed in

Waugh et al. (1999), the variations in dynamics (i.e., eddy

wave driving) do not explain the longer term trends of the

breakup in the SH. Hence, the observed late breakups in

the SH (and possibly the NH) are presumably related to

radiative changes from ozone losses.

3.2 BASIC POLAR STRATOSPHERIC

PROCESSES

3.2.1 Transport and Dynamics

This section discusses the structure and dynamics

of the polar stratosphere, including trace gas transport.

Section 3.2.1.1 gives a brief overview of the mean vortex

structure, providing a context for the later assessment.

Section 3.2.1.2 assesses recent studies of the dynamics

and structure, including variability and trends. An important point is the apparent delay in polar vortex breakdown

in springtime of the Southern (and possibly Northern)

Hemispheres; extending the cold winter season is a critical factor in increasing the likelihood of chemical ozone

loss, but even in the absence of chemical processes, the

continued isolation of the polar region into the springtime

leads to a “dynamical” ozone deficit. Section 3.2.1.3

examines transport processes, including the mean meridional circulation and issues of transport inside, outside,

and across the boundary of the polar vortex; these play a

crucial role in fixing the distributions of ozone and other

trace gases, which impacts both the physical processes

(such as radiative heating) and chemical ozone loss.

3.2.1.1 THE POLAR VORTEX: MEAN STRUCTURE

The winter stratospheric circulation is dominated

by the polar night jet, which is at the edge of the polar

vortex. Understanding the polar vortex dynamics is central to our ability to understand recent ozone change and

to predict future ozone.

The polar vortex structure is well understood:

absence of solar heating in winter leads to low temperatures, which are offset by the adiabatic warming caused

by the descending branch of the Brewer-Dobson circulation. This circulation is caused by the damping of planetary and gravity waves in the middle atmosphere (e.g.,

Fels, 1985). An illustration of the polar vortex (Figure 3-

11) shows the polar night jet (peaking near 60°N at about

45 km) and the strong descent (shown by the meridional

stream function). Descent in the polar region leads to (diabatic) downward transport, carrying ozone and other trace

gases from the mesosphere to the lower stratosphere

during winter (e.g., Rosenfield and Schoeberl, 2001).

The asymmetry of the polar vortices in the two

hemispheres is a consequence of the different topographic

features: the weaker wave activity propagating from the

SH troposphere provides less forcing and therefore a

weaker Brewer-Dobson circulation than in the NH (e.g.,

Randel and Newman, 1998). The Antarctic vortex is more

symmetric, stronger, and colder than the Arctic, as illustrated by the 50-hPa geopotential height distributions in

middle winter (Figure 3-12). Even in the absence of chemical ozone destruction, these dynamical differences lead

to substantially more ozone in the Arctic than in the

Antarctic vortex, especially in springtime.

The main impacts of these dynamical differences

on Arctic (compared with Antarctic) ozone arise from the

stronger diabatic descent, which transports trace species

downward more rapidly, the weaker isolation of the Arctic,

and the decreased likelihood of PSC formation. Since

PSCs form at temperatures near 195 K at 50 hPa (see

Section 3.2.2), they can form every winter in the Antarctic

vortex core, but only on colder-than-average days in the

Arctic (e.g., Pawson et al., 1995; Pawson and Naujokat,

1999).

While these basic mechanisms that determine the

vortex structure and tracer transport in polar regions are

now well understood, there are important aspects for

which the complexity is only partially described. Recent

results pertaining to these uncertainties are assessed in the

next section (3.2.1.2).

3.2.1.2 POLAR VORTEX: CAUSES OF INTERANNUAL

VARIABILITY AND ITS IMPLICATIONS

Differences between the hemispheres, caused by

the stronger wave driving in the NH, are also evident in the

year-to-year variations. Figure 3-9 shows what appear to

be robust temperature trends in the springtime, but interpretation of such trends is complex, because of the large

interannual variability of high-latitude temperature (Figure

3-8), which varies with time of year and is different in the

two hemispheres (Figure 3-13; e.g., Scaife et al., 2000b).

Southern Hemispheric variability peaks in late winter and

spring (e.g., Kuroda and Kodera, 1998), whereas variability in the NH is large throughout the season (e.g.,

Labitzke, 1982). The variability of the stratosphere

(defined in terms of departures from the long-term mean)

is characterized by a “see-saw” of temperature and mass
POLAR OZONE

3.14

between the polar region and midlatitudes (e.g., Labitzke,

1982; Kodera et al., 1996): anomalously weak wave

forcing leads to a strong polar vortex and a weak BrewerDobson circulation, with a cold polar region and warmer

midlatitudes (the converse is true for strong wave forcing).

Newman et al. (2001) demonstrated the quantitative

linkage between the upward-propagating wave activity

through the tropopause region and the strength of the polar

vortex.

The occurrence of a strong, cold polar vortex leads

to anomalously low ozone in the polar region, because the

transport of ozone-rich air is weak and because the potential for PSC processing, a precursor to chemical ozone

loss, is enhanced. This means that in years with weak tropospheric wave forcing, a stronger polar vortex will result

in less ozone in the polar region. Chipperfield (1999)

reported results from a 6-year simulation using the

SLIMCAT chemical transport model (CTM), driven by

the United Kingdom Meteorological Office (UKMO)

analyses of the meteorology. The horizontal winds and

temperatures are taken from the UKMO analyses, and the

vertical motion is diagnosed using a radiation scheme.

The model simulates the interannual variations in chlorine activation during northern winters and reproduces

the repeatable pattern of activation observed during

southern winters. Chipperfield and Jones (1999) utilized

the same model to evaluate the relative contributions of

photochemical and dynamical processes to interannual

-90 -60 -30 0

Latitude

0

10

20

30

40

50 Altitude (km)

J

J

0 30 60 90

Latitude

1000

100

10 Pressure (hPa)

J

J

-20 0 20 40 60 80 100

m s

July January

Figure 3-11. Zonal-mean sections of zonal velocity (shaded), the mean meridional streamfunction (black

lines), and an atmospheric trace gas (white contours) for midwinter (July in the Southern Hemisphere and

January in the Northern Hemisphere). The data were obtained from multiannual climatologies. The figure

illustrates the stronger polar night jet in the Southern Hemisphere, the stronger subtropical jet in the northern

winter, along with consistent differences in the strength and structure of the mean meridional circulation and

trace gas distributions.
variability in northern high-latitude ozone, and showed

that dynamical variations dominate interannual variability.

Hadjinicolaou et al. (2002) find similar results using a

long run of the same transport model, driven by European

Centre for Medium-Range Weather Forecasts (ECMWF)

analyses but using a simplified chemical scheme.

In the NH, the anomalies in polar vortex strength

are a part of what is now known as theArctic Oscillation

(AO) (Thompson and Wallace, 1998). The AO and its

Southern Hemispheric counterpart are also referred to as

the annular modes. The annular structure of the AO in the

stratosphere can be traced to the surface, with a strong

link to the North Atlantic Oscillation (NAO) in the NH

(Thompson and Wallace, 1998). Whereas there is some

debate about the role of the stratosphere in forcing anomalies in the tropospheric component of the AO (e.g.,

Perlwitz and Graf, 2001; Ambaum et al., 2001), that is

beyond the scope of this Assessment of stratospheric

ozone. This discussion focuses on the stratospheric

component of the AO.

A high (low) AO index corresponds to a strong

(weak) vortex and low (high) polar ozone column values

(Thompson and Wallace, 2000; Hartmann et al., 2000).

Thompson et al. (2000) estimated that approximately 40%

of recent apparent polar ozone loss in March could be

explained by the tendency of the AO to remain positive in

the springtime, which describes a strong, cold, and isolated polar vortex. However, on the basis of the observations alone, it cannot be determined whether the signal in

ozone is caused by the AO anomaly, or whether the AO

anomaly is a consequence of ozone depletion, or whether

both are coherently forced by some other factor.

There is strong evidence that the AO signal originates near the subtropical stratopause and propagates

poleward and downward through the mechanism of

wave forcing (e.g., Baldwin and Dunkerton, 1999,

2001; Kuroda and Kodera, 1999; Kodera et al., 2000;

Christiansen, 2001). Kodera and Kuroda (2000) show

how the interannual variability of wave forcing can cause

such anomalies to take different phases in different years.

Isolating causes of variability and the factors that

drive trends is not straightforward. Apart from the link

between the AO strength and the upward propagation of

planetary waves, other mechanisms have been related to

the polar vortex. Model simulations reveal that a substantial year-to-year variability in the stratospheric vortex

(and, hence, the AO) can exist in the absence of variations

in boundary conditions or other forcing mechanisms (e.g.,

Yoden et al., 1999; Hamilton et al., 1999). This variability,

forced by internal dynamics of the atmosphere, means that

POLAR OZONE

3.15

Geopotential Height (gpdm) and Temperature (K) at 50 hPa

January July

210

Figure 3-12. Polar stereographic projections for January in the Northern Hemisphere and July in the Southern

Hemisphere. The geopotential height (red contours; geopotential decameters (gpdm)) and temperature (black

contours; K) for 50 hPa are shown. Light-blue shading indicates regions colder than 200 K, while dark-blue

shading shows regions colder than 195 K. Data are 19-year means from the United Kingdom Meteorological

Office-processed TOVS data (Scaife et al., 2000b).
POLAR OZONE

3.16

many factors often invoked as causes of interannual variability in the real atmosphere may or may not be significant. Despite this, there is some evidence of coupling

between the polar vortex and other atmospheric variations; the main relationships that have been studied are

the QBO of tropical winds, the 11-year variability of solar

radiation, the phase of the ENSO, and major volcanic

eruptions. Although polar vortex composites grouped

according to these mechanisms show apparent signals,

several factors complicate their interpretation and robustness. The most severe complications are that the observational record covers only about four decades and that

some of the forcing factors vary in unison. For instance,

following Labitzke and van Loon (1997) and grouping

northern midwinter polar vortex structure according to

the solar cycle (high or low) and the phase of the QBO

(East or West) leads to 12 winters in the low/West category, with a strong polar vortex; however, 5 of these 12

winters coincide with ENSO cold events or volcanic eruptions, which have the same anomalies (see Figure 3-14).

Determining robust relationships from observations on

the basis of these overlapping factors and the internal

variability is thus impossible.

Models have been used to address these questions.

The SKYHI general circulation model (GCM) with an

artificially forced QBO reproduces observed QBO-related

interannual variability in the Arctic vortex (stronger when

the tropical winds are westerly, weaker when easterly)

and variations in wintertime stationary wave patterns

(Hamilton, 1998). Shindell et al. (1999b) found that the

QBO significantly modulated the strength and propagation of planetary-wave energy in the troposphere in the

Goddard Institute for Space Studies (GISS) model,

leading to higher (3- to 5-K) zonal-mean temperatures at

high southern latitudes for late winter and early spring

during the QBO easterly phase. Niwano and Takahashi

50 hPa, 50-90oS Minimum Temperature

180

190

200

210

220

230

K

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec 1978-2001 2001

Type I PSC

HNO = 6 ppbv, H O = 4 ppmv

Type II PSC

50 hPa, 50-90oN Minimum Temperature

180

190

200

210

220

230

K

Jul Aug Sep Oct Nov Dec Jan Feb Mar Apr May Jun

1978-2001 1999-2000

Type I PSC

HNO = 6 ppbv, H O = 4 ppmv

Type II PSC

Figure 3-13. Time series showing

binned distributions of 50-hPa minimum

polar temperatures for 50°-90°N (top)

and 50°-90°S (bottom). The blue line

shows the 1978-2001 mean, and the

thin black lines show the maximumminimum values. Shading shows the

density of observations, with heavy

shading indicating a high probability and

light shading indicating a low probability.

Data are binned into intervals of 5 days

by 2 K, and therefore the maximum

number of observations in any bin is 95

(for 19 years of data). The green line

shows the values for 1999-2000 (top

panel) and 2001 (bottom panel). The

seasonal cycles are offset by 6 months

so that the seasonal progression in each

hemisphere is contrasted. Thresholds for

PSC formation are indicated by the horizontal lines. Data were processed as in

Scaife et al. (2000b).
POLAR OZONE

3.17

(1998) studied the influence of the QBO on the NH winter

circulation; their model reproduced the relationship

between the polar vortex strength and the QBO phase and

a related NAO pattern in the troposphere. These and earlier studies have worked on the premise that the lower

stratospheric winds impact planetary wave propagation,

whereas more recent work (Gray et al., 2001) has shown

that the polar vortex anomalies are more strongly related

to winds near the tropical stratopause (which are indirectly

affected by the QBO).

There are two important factors in isolating the

impacts of solar forcing on the circulation and climate.

First, stratospheric ozone changes modulate the response

of the temperature to the changes in solar irradiance

Figure 3-14. The anomalies in January/February 30-hPa geopotential height revealed by composite anomaly

plots for ENSO and volcanic events, 1958-1997, as determined from Freie Universität Berlin analyses. The

two top panels are (a) the composite anomaly of the seven winters with ENSO cold events, and (b) that for the

nine warm events. The anomalies are in gpdm and are relative to the mean of the years with no ENSO events.

The lower two panels show the warm-event anomalies of panel (b), but further split into (c) the 6 years with no

volcanic events and (d) the 3 years with them. These results were updated from van Loon and Labitzke (1987)

and Labitzke and van Loon (1989).
POLAR OZONE

3.18

(Haigh, 1994). However, studies of solar impacts on

ozone, mostly using two-dimensional (2-D) models, have

been plagued by an inability to reproduce the solar-ozone

relationship detected in observations (e.g., Brasseur, 1993;

Hood and Zhou, 1999). Second, inclusion of the correct

spectral dependence of solar irradiance variations in the

atmospheric heating rate calculations is essential to capture the correct vertical structure of heating rates (Haigh,

1999; Shindell et al., 1999a; Larkin et al., 2000). However, even incorporating these feedbacks, climate model

studies generally remain inconclusive about the role of

solar-induced perturbations in the variability of the Arctic

polar vortex.

Volcanic aerosol loading can possibly impact polar

ozone by perturbing stratospheric chemistry and transport. Chemical perturbations from heterogeneous reactions on aerosols are discussed in Section 3.2.2. Increased

volcanic aerosol loading of the tropical lower stratosphere

leads to a warmer tropical lower stratosphere some months

after a volcanic eruption (e.g., Robock, 2000), which is

discussed in more detail in Chapter 4. The polar response

to this tropical warming has a northern polar vortex

remaining anomalously cool in the winter following the

eruption (e.g., Kodera, 1994), as shown in Figure 3-14.

Figure 3-14 shows that the three volcanic eruptions

affected winters with ENSO warm events and that they

“reverse” the ENSO anomalies, leading to anomalously

strong, cold polar vortices (van Loon and Labitzke, 1987;

Labitzke and van Loon, 1989), thereby increasing the likelihood of negative polar ozone anomalies. Based on only

three events (and because of the factors discussed above),

these results must be interpreted with caution. This relationship has also been isolated in models (Kirchner et al.,

1999; Ramachandran et al., 2000).

Although temperatures low enough for PSC formation occur on the large scales (e.g., Pawson and

Naujokat, 1997, 1999), the likelihood of their occurrence

is enhanced by the temperature perturbations induced by

medium-scale waves (e.g., Grewe and Dameris, 1997;

Sato et al., 2000; Teitelbaum et al., 2001), as well as by

inertial gravity waves (e.g., Dörnbrack et al., 2001, 2002).

The temperature perturbations induced by these waves

can cause sufficient additional cooling for PSCs to form

in locations where the large-scale flow would not support

them. This is particularly important on the vortex edge,

where the processed air can be irreversibly transported

into the middle latitudes (in the presence of breaking

waves) and where the air masses are more likely to be illuminated, enhancing the potential for ozone depletion.

Ozone mini-holes occur because of synoptic-scale,

reversible advection (e.g., McKenna et al., 1989; Newman

et al., 1988) related to upper tropospheric anticyclonic

structures. The high tropopause, coupled with ascending

motion, leads to extremely low total ozone values with

lifetimes of up to several days. Steinbrecht et al. (1998)

show that correlations in tropopause height correlate with

ozone concentration changes in the region up to 23 km,

illustrating the depth of the disturbances. However, the

low ozone values themselves are short-lived features that

are unrelated to chemical loss. The dynamical forcing

that causes ozone mini-holes also causes adiabatic

cooling, which can lead to synoptic-scale temperature perturbations of sufficient magnitude to allow PSC formation (e.g., McKenna et al., 1989; Grewe and Dameris,

1997). The importance of baroclinic disturbances in producing PSC formation near the polar vortex edge has been

discussed by Hood et al. (2001) and Teitelbaum et al.

(2001). Orsolini and Limpasuvan (2001) showed how

these disturbances are linked to the storm tracks, which

vary in unison with the AO. There is thus a flowdependent nature to the likelihood of synoptic-scale PSC

formation and to the likelihood that PSCs contribute to

ozone loss on the vortex edge region.

Just as baroclinic waves help PSC formation on the

synoptic scales, mesoscale disturbances from gravity

waves are also important. Volkert and Intes (1992)

demonstrated PSC formation in wave crests over

Scandinavia in their model of topographically forced

gravity waves. The importance of gravity-wave PSCs was

also demonstrated by Deshler et al. (1994) and Meilinger

et al. (1995). High-resolution radiosonde data have

recently provided much-needed information on stratospheric gravity-wave morphologies in and around the

Antarctic (Pfenniger et al., 1999; Zink and Vincent, 2001)

and Arctic (Whiteway and Duck, 1999; Yoshiki and Sato,

2000). Although the microphysical effects of a background spectrum of gravity waves are smaller than first

thought (Bacmeister et al., 1999), it is now accepted that

mesoscale temperature decreases due to large-amplitude

gravity waves, particularly mountain waves, can lead to

temperatures low enough for PSC formation (Carslaw et

al., 1998b, 1999; Schulz et al., 2001), and lead to structure inside larger scale PSCs (Toon et al., 2000). PSCs

are discussed in Section 3.2.2 of this chapter.

Important advances have been made in our ability

to model lower stratospheric gravity waves and, especially, to resolve such waves in global meteorological

analyses. Dörnbrack et al. (2001) demonstrated that the

high-resolution ECMWF operational meteorological

analyses capture gravity-wave structures over Scandnavia. This represents an important advance for applying

the analyses to our understanding of the gravity-wave

morphology and its importance for PSC formation. An

additional advance of some importance was made by
POLAR OZONE

3.19

Dörnbrack et al. (2002), who detected inertia-gravity

waves over Scandinavia in the ECMWF analyses and in

situ data, noting their role for PSC formation. The isolation of gravity waves in such operational analyses points

to their potential utility in mountain wave forecasting and

analysis, meaning that the off-line models that have been

used for such studies could eventually become unnecessary.

Another important role played by gravity waves

(from all sources) is that they transport momentum into

the middle atmosphere; as these waves break, they deposit

momentum to the mean flow, constituting an important

driving mechanism for the Brewer-Dobson circulation.

The importance of these waves for driving the flow and

reducing biases in global models is discussed in more

detail in Section 3.5.2.1.

3.2.1.3 POLAR TRANSPORT AND MIXING

This section discusses in detail the physical

processes that lead to the redistribution of trace gases in

the polar regions. Trace gas distributions are determined

by the balance between the slow, mean-meridional circulation and the more rapid, quasi-isentropic mixing (e.g.,

Holton, 1986); the following discussion examines these

components of transport in and around the polar vortex

and the exchange across the vortex edge.

Manney et al. (2002) examine impacts of using different meteorological analyses, which affect the amount

of exchange between middle latitudes and the vortex.

Despite the uncertainties, the consensus is that the vortex

remains quite isolated in wintertime. Vertical transport

leads to descent of tracers in and around the vortices, while

mixing redistributes ozone and trace gases on isentropic

levels; there is some exchange across the vortex edge,

associated with large-scale mixing events.

Descent in the Polar Vortex

Descent inside the polar vortex builds up (or maintains) lower stratospheric ozone over the winter, making

it an important process to understand. The descending

branch of the Brewer-Dobson circulation is driven by

wave forcing of the flow (see Section 3.2.1.1). Descent

rates can be determined in several manners: (1) “directly”

from the vertical velocities produced in routine meteorological analysis systems, such as the UKMO (Swinbank

and O’Neill, 1994) and the Data Assimilation Office

(DAO) (Rood et al., 1997), (2) based on the crossisentropic transport determined by diabatic heating rates,

and (3) using measurements of long-lived trace gases with

well-understood vertical gradients. The various estimates

are in reasonably good agreement, showing stronger

descent in the upper stratosphere than in the lower stratosphere and unmixed descent from the upper to the lower

stratosphere. The strongest unmixed descent occurs in

the Antarctic vortex. Descent in the Arctic vortex is more

variable; it occurs on the vortex edge when the temperature is higher there (e.g., Manney et al., 1999).

Schoeberl et al. (1995) used Halogen Occultation

Experiment (HALOE) methane (CH4) data to estimate

the descent rate as 1.8 km/month inside the Antarctic

vortex in February to October 1992 (see Appendix 3A for

HALOE data description). Descent rates over Antarctica

were deduced from the Improved Stratospheric and

Mesospheric Sounder (ISAMS) carbon monoxide (CO)

data for April to July 1992 by Allen et al. (2000) and from

the Polar Ozone and Aerosol Measurement (POAM) III

H2O data by Nedoluha et al. (2000). Abrams et al. (1996)

demonstrated strong descent in the upper stratosphere

(3.2 km/month at 40 km) with weak descent in the lower

stratosphere (0.8 km/month at 20 km) based upon the

Atmospheric Trace Molecule Spectroscopy (ATMOS)

data from November 1994. Kawamoto and Shiotani

(2000) also used HALOE data and UKMO meteorological analyses to investigate the interannual variability of

the descent rate, using the February-October (winter)

averages in 1992 to 1997 HALOE CH4 data (see Appendix

3A for POAM and HALOE data descriptions). They

found that the descent varies between 1.2 and 1.8

km/month using the 0.6 parts per million by volume

(ppmv) CH4 contour inside the polar vortex and is consistent with the wave driving determined from the UKMO

analyses. To summarize, the various estimates of polar

descent rates give reasonably consistent results, given that

they are presented for different levels, seasons, and years.

The Vortex Core

The degree of mixing within the vortex core has

come under scrutiny because of assumptions about representing the bulk behavior of the vortex with irregular temporal and spatial sampling. If the core is well mixed, then

measurements anywhere within the vortex will suffice to

characterize its behavior. If the core is not well mixed,

then more frequent sampling at separated locations is necessary to do this. Schoeberl et al. (1990) assumed that the

vortex was relatively well mixed in assessing ozone loss

using Earth Resources-2 (ER-2) data from the Airborne

Arctic Stratospheric Expedition (AASE)-I mission during

the Arctic winter of 1988/1989. Richard et al. (2001) used

ER-2 data from the SOLVE-THESEO 2000 campaign to

show that tracer-tracer relationships inside the vortex

during the Arctic winter of 1999/2000 are distinct and

compact, suggesting a rapid mixing in the Arctic vortex.

Whereas current evidence suggests that the Arctic vortex
is relatively well mixed in the absence of intrusions of air

from the vortex edge, Lee et al. (2001) present evidence

that the Antarctic vortex is separated into two regions: a

strongly mixed vortex core and a weakly mixed ring of

air extending to the vortex boundary. 

Transport Across the Vortex Edge and Mixing

The balance of mass and trace gases in and around

the polar vortex is determined by the downward transport

and exchange across the vortex edge. Any vertical gradient in the vortex-averaged mass flux will be compensated for by flow across the vortex edge. The discussion

below will separate the transport across the vortex edge

in winter from transport occurring as the polar vortex

breaks down.

Planetary-wave breaking can be responsible for

vortex shrinking as well as sharpening of the vortex edge

(e.g., Thuburn and Lagneau, 1999). A number of studies

have shown considerable variability in the width of the

Antarctic vortex edge when it is perturbed (Teitelbaum et

al., 1999; Perez et al., 2000), which results in nonlinear

irreversible transport and mixing of vortex air into midlatitudes (Teitelbaum et al., 1999). Lidar observations at

Dumont d’Urville (66.4°S, 140°W) allow sampling at and

around the moving vortex edge (Godin et al., 2001); observations of Mt. Pinatubo aerosols, made during October

and November 1992, show the sharpness of the vortex

edge and low mixing between the inner vortex and the

outside air above 400 K.

The first obstacle to determining the cross-vortex

flow is to unambiguously define the vortex edge. Chen

(1994) defined it as the potential vorticity (PV) contour

that has the smallest lengthening rate; he found a vertical

dependence to the transport across the edge, with more

transport out of the vortex at potential temperatures lower

than 400 K than at higher levels. Tuck et al. (1995)

reached similar conclusions using ER-2 data, showing

also that the vortex edge region can be quite wide. This

viewpoint of the polar vortex as a reasonably well isolated

entity is now generally accepted, but the amount of

“leakage” from the vortex (as a function of altitude) is not

yet well understood. Different proposed definitions of

the vortex edge include the wind maximum and the

strongest gradients in PV (e.g., Bowman, 1996; Nash et

al., 1996). The uncertainty in defining the vortex edge

remains, so there is no unambiguous estimate of the vertical structure of cross-vortex transport. Mechanisms for

the transport are at least qualitatively understood and are

discussed here.

Recent results continue to sustain our understanding of the vortex edge impermeability (Chen, 1994)

and of the polar vortex as a quasi-isolated containment

vessel. Norton and Chipperfield (1995) and Jones and

MacKenzie (1995) had argued that ozone-depleted air

from the polar vortices makes only a small contribution

to middle-latitude ozone loss. High-resolution, singlelevel models with weak dissipation (e.g., Juckes and

McIntyre, 1987; Mo et al., 1998; Thuburn and Lagneau,

1999; Sobel and Plumb, 1999) have further confirmed that

the export of air from the polar vortex is constrained.

Vincent and Tranchant (1999) also found little mixing

across the vortex edge at 520 K in the Antarctic. Li et al.

(2002) used a CTM driven by analyzed winds to show that

less air is indeed exported from the Antarctic polar vortex

to middle latitudes than descends into the troposphere.

Laminae and filaments could be an important

mechanism in the mixing of air across the quasiimpermeable vortex edge. Such structures are common

in winter and spring. As pointed out in the previous

Assessment (WMO, 1999) the filaments/laminae, i.e.,

material sheets that tilt outward with increasing height

(Schoeberl and Newman, 1995; Newman and Schoeberl,

1995) with initial horizontal scales of a few thousand kilometers, can lead to irreversible mixing on the time scale

of 20 to 25 days over which they decay. Laminae have

been detected in a variety of data types, including in situ

aircraft observations (Newman et al., 1996); sondes and

lidar (Bird et al., 1997; Orsolini et al., 1997; Teitelbaum

et al., 2000); and satellites (Manney et al., 1998, 2000).

They have been successfully modeled (e.g., Orsolini et

al., 1997). Waugh and Dritschel (1999) analyzed the relationship between Rossby wave breaking and vortex structure. These studies show that filamentation can lead to

vortex air being peeled off and eventually mixed irreversibly into the surf zone, although some air may rejoin

the polar vortex. Furthermore Manney et al. (1998, 2001)

have also shown that lamination processes within the polar

vortex did not result from exchange across the vortex edge

but rather from transport variations within the vortex.

Hence, there is a need to assess the behavior of

laminae and the magnitude of their contribution to the

total exchange between the vortex and midlatitudes.

Appenzeller and Holton (1997) attempted to diagnose the

production of tracer laminae using satellite data and meteorological analyses, as a first step in determining their

contribution to transport. Nevertheless, there arose some

limitations regarding the use of such a diagnostic

(Kettleborough and Holton, 1999), because (1) it does not

include small vertical scales that are relevant in defining

tracer lamination, and (2) there can be reversible contributions. In other words, the proposed diagnostic could

overestimate the transport and mixing.

A principal challenge to modeling polar ozone in

the NH and the effects of polar processes on middle latiPOLAR OZONE

3.20
POLAR OZONE

3.21

tudes is to ensure that the models produce the appropriate

balance among the many processes that contribute directly

or indirectly to the polar lower stratospheric ozone tendency. Both transport and photochemical processes contribute. The year-to-year variability in meteorological

fields is significant, and the northern vortex may be cold

and strong, as in the 1996/1997 and 1999/2000 winters,

or warmer and more disturbed, as in the 1997/1998 winter

(Sinnhuber et al., 2000; Guirlet et al., 2000). Simulations

have focused on replicating observations for ozone and

other trace gases, and quantifying model sensitivity to

various processes (Chipperfield and Pyle, 1998). Such

studies point out the importance of developing a better

understanding of the physical processes (e.g., those

leading to denitrification), so that model parameterizations respond appropriately to changes in temperature,

water vapor, or nitric acid (HNO3) that may result from

climate change.

Millard et al. (2002) utilized diagnostics developed

by Lee et al. (2001) to quantify the importance of polar

processes to ozone change at middle latitudes. Both the

chemical processes that contribute to polar ozone loss and

the transport processes that impact mixing between high

latitudes and middle latitudes vary depending on the

meteorology of a particular year. CTMs have been utilized to address questions concerning specific winters.

For example, Lefèvre et al. (1998) utilized a CTM forced

by winds from ECMWF to show that both transport and

photochemical processes contributed to the record-low

ozone observed by TOMS during northern spring 1997

(Figure 3-2). Like the studies of Guirlet et al. (2000), their

results are broadly consistent with observations.

However, as shown by Douglass et al. (2001), the model

results separating photochemical and transport contributions are sensitive to the vertical velocity and to the ozone

vertical gradient.

The breakdown of the polar vortex, whether by a

major midwinter warming or in the final warming, allows

vortex air to be mixed relatively easily with air from

middle latitudes. Atkinson and Plumb (1997) showed that

as the Antarctic vortex breaks down, a substantial amount

of ozone-depleted air is transported to middle latitudes.

Once there, it can effectively mix with the ambient air

masses. Effective diffusivity has been used as a diagnostic

for mixing by Allen and Nakamura (2001), who show

increases in mixing lengths as the polar vortices break up.

3.2.2 Polar Stratospheric Clouds

PSCs play two important roles in polar ozone

chemistry. First, the particles support chemical reactions

leading to active chlorine formation, which can catalytically destroy ozone. Second, nitric acid (HNO3) removal

from the gas phase can increase ozone loss by perturbing

the reactive chlorine and nitrogen chemical cycles in late

winter and early spring.

PSCs are divided into two main categories. Type I

PSC particles contain nitric acid, either in the form of

liquid ternary solutions with water and sulfuric acid or as

solid hydrates of nitric acid. Type II PSCs are made of

ice particles. Knowledge of PSC particle sizes, number

concentrations, composition, phase, and evolution is central to efforts to develop prognostic models of how PSCs

affect the chemistry of the polar stratosphere. In situ

observations of PSCs from balloons or aircraft are often

used to obtain detailed information on cloud particle size

distribution and composition. Remote sensing platforms,

such as lidar and satellites, provide complementary information on phase and large-scale time evolution of PSCs,

respectively. We now briefly review recent advances in

our understanding of PSC properties and their effect on

denitrification and dehydration.

3.2.2.1 OBSERVATIONS OF PSC PHYSICAL

PROPERTIES AND THEIR INTERPRETATION

In Situ

The previous Assessment (WMO, 1999) described

considerable improvements in our understanding of

liquid PSCs but highlighted the outstanding uncertainties in the properties of solid particles. Solid nitric-acidcontaining PSC particles are important because, in contrast to the submicron liquid aerosol, they may be present

with sufficiently low number concentrations (<10-2 cm-3)

to allow a few particles to grow to large sizes, leading to

sedimentation and denitrification. Our understanding of

the range of solid-particle number concentrations and

sizes that can form in the polar stratosphere has improved

since the previous Assessment as a result of new in situ

observations.

Observations in the Arctic stratosphere at altitudes

from 16 to 20 km in January to March 2000 detected a

population of large nitric acid particles with very low

number concentrations (Fahey et al., 2001; Northway et

al., 2002a); see Figure 3-15. Large particles, with sizes

and number concentrations similar to those observed by

Fahey et al. (2001), were detected by the Multiangle

Aerosol Spectrometer Probe (MASP) (Carslaw et al.,

2002). These measurements (Fahey et al., 2001;

Northway et al., 2002a) are very important because they

provide conclusive evidence that such large particles are

composed principally of nitric acid (probably present as

nitric acid hydrates).
Observations of large nitric acid particles raise several questions. The most obvious question is how these

particles compare with previous observations in the Arctic

and Antarctic. Balloonborne instruments such as the

optical particle counter (OPC) (e.g., Deshler et al., 1991,

1994; Deshler and Oltmans, 1998; see also WMO, 1999,

and references therein), which has flown in many previous

Arctic winters, are capable of detecting particles up to 20

mm in diameter with number concentrations greater than

about 6 ¥ 10-3 cm-3. However, this is higher than the

average particle concentrations (~10-4 cm-3) measured by

Fahey et al. (2001). It is interesting to note that in earlier

Antarctic measurements, the OPC instrument, which operated at a different inlet flow rate, detected large PSC particles with number concentrations as low as 10-4 cm-3

(Table 3-1). The forward scattering spectrometer probe

(FSSP) instrument, which measured PSC size distributions during the 1989/1990 Arctic winter (Dye et al.,

1992), detected PSCs with total number concentrations of

about 10-3 cm-3. Some particles were observed up to 13-

mm diameter with concentrations of ~few ¥ 10-4 cm-3

mm-1, similar to those observed by Fahey et al. (2001).

These large particles were attributed to ice (Wofsy et al.,

1990), although some of them could certainly have been

nitric acid trihydrate (NAT). Thus populations of very

few (~10-4 cm-3) large PSC particles have been observed

previously, but it is only through recent measurements

(Fahey et al., 2001) that we have learned that such large

particles are indeed enriched in nitric acid.

The second question is whether these large particles present at very low number concentrations can denitrify the stratosphere. Simple calculations presented by

Fahey et al. (2001) demonstrate that large nitric acid particles can grow to their optimal observed sizes in about 5

to 8 days, implying that they must have nucleated several

kilometers above the aircraft flight altitude. Instantaneous

downward flux calculations of nitric acid contained in

such large particles indicate that the large particles were

capable of causing significant denitrification (Fahey et

al., 2001; Northway et al., 2002b). Three-dimensional

model simulations of particle growth further show that

the observed PSC sizes are consistent with growing NAT

and/or nitric acid dihydrate (NAD) particles (Carslaw et

al., 2002). However, it is important to note that PSCs with

size distributions differing from those measured in winter

1999/2000 have been observed in previous winters in both

hemispheres (see Table 3-1). Model simulations by Jensen

et al. (2002) show that previously observed PSC particle

size distributions (Dye et al., 1992; Hofmann and Deshler,

POLAR OZONE

3.22

10-2

2

3

4

5

6

10-1

2

3 Equiv. mixing ratio (ppbv)

7

108

2

3

4

5

6

7

109 Flux (molecules cm-3 km day-1)

5 10 15 20

Particle diameter (μm)

10-6

10-5

10-4

10-3 Concentration (# cm-3 μm-1)

A

B

C

Figure 3-15. Results from the statistical simulation of total reactive

nitrogen (NOy) values observed by the ER-2 NOy instrument on 20

January 2000 (Fahey et al., 2001). Panels show (A) the adjusted

best-fit size distribution and number concentration, (B) the equivalent gas-phase mixing ratio of HNO3 derived from the distribution,

and (C) the derived flux of HNO3 at 60 hPa (~19.5 km). The size

distribution is given by the Gaussian functions in the caption of

Figure 3 of Fahey et al. (2001) adjusted to account for sizedependent sampling efficiency. The adjustments, which are in addition to a basic particle enhancement factor of 12.8, range from +30%

to –10% over the size range and are calculated with a fluid dynamical model of the particle separator. The concentration integral of

the large (small) mode is 2.3 ¥ 10-4 cm-3 (2 ¥ 10-3 cm-3) with an estimated uncertainty of ±30%. The lower limit of the vertical axis in

(A) corresponds to the detection of a single particle over the 800-s

observation period. The total gas-phase HNO3 of the large (small)

mode is 1.5 ppbv (0.2 ppbv). The HNO3 flux of the combined distribution is 5 ¥ 109 molecules cm-3 km day-1 or 2.2 ppbv km day-1 at 60

hPa. Adapted from Fahey et al. (2001).
POLAR OZONE

3.23

1991), with number concentrations in the range of 10-2 to

10-3 cm-3, are also capable of efficiently denitrifying the

polar stratosphere.

A third question is how these large nitric acid particles form in the polar stratosphere. Both homogeneous

(Tabazadeh et al., 2001) and heterogeneous (Tolbert and

Toon, 2001; Drdla et al., 2002) freezing mechanisms have

been suggested to account for the formation of large nitric

acid particles. For such nucleation mechanisms to

operate, the cooling caused by synoptic-scale uplift of air

masses (Teitelbaum et al., 2001; Spang et al., 2001;

Hendricks et al., 2001; Saitoh et al., 2002) can provide

favorable conditions for solid PSCs to form. Laboratory

observations show that concentrated aqueous nitric acid

aerosols can homogeneously crystallize into hydrates of

nitric acid (Disselkamp et al., 1996; Bertram and Sloan,

1998a, b; Prenni et al., 1998; Salcedo et al., 2001). The

stratospheric particle system has also been studied using

thin films, where gas-phase nitric acid and H2O are

absorbed by cold aqueous sulfuric acid solutions (Iraci et

al., 1994, 1995, 1998). The results of these thin-film

experiments show that HNO3 uptake in sulfuric acid can

cause freezing of nitric acid hydrates in solution.

Tabazadeh et al. (2001) have recently extrapolated the

laboratory homogeneous freezing rates of Salcedo et al.

(2001) and obtained nucleation rates sufficient to produce

large nitric acid particles in a microphysical model. Note

that the nucleation rates extrapolated from Salcedo et al.

(2001) are much higher than upper limits derived from

earlier bulk freezing experiments with 1-milliliter samples, which used stratospheric temperatures and liquidphase compositions (Koop et al., 1995, 1997). Modeling

studies (Drdla et al., 2002) further show that heterogeneous freezing can also produce large particles if only a

very small fraction (<0.1%) of stratospheric aerosol particles contained an effective freezing nucleus, although it is

not clear what the freezing nucleus should be (Biermann

et al., 1996). In addition to direct homogeneous and heterogeneous freezing mechanisms, the large particles may

also form by a gradual sedimentation from the base of

“mother clouds” containing much higher number concentrations of small solid particles (Füglistaler et al., 2002;

Dhaniyala et al., 2002), such as those generated by lee

wave clouds (Carslaw et al., 1998a).

The freezing mechanisms described above are all

capable of producing low number concentrations of

sedimenting nitric acid particles. However, model simulations do strongly suggest that the large nitric acid particles, observed during the winter of 1999/2000, were

unlikely to have nucleated in synoptic-scale ice clouds,

which were not sufficiently prevalent in the days preceding the observations (Carslaw et al., 2002; Drdla et al.,

2002). Further analysis of satellite data in the Antarctic

also seems to suggest that large nitric acid particles formed

in the winter of 1992 independent of synoptic-scale ice

clouds (Tabazadeh et al., 2000). Formation of denitrifying particles when polar temperatures are above the ice

frost point will allow denitrification to occur at higher

temperatures than previously assumed in 3-D chemical

models (see Section 3.2.2.3).

A final question is whether these large particles

observed in situ can also be detected by lidar. Aircraft

lidar observations in January to March 2000 detected

regions of enhanced aerosol backscatter in regions where

large nitric acid particles were detected (Flentje et al.,

2000). However, Flentje et al. (2000) inferred an approxTable 3-1. Observations of solid PSC number density and particle size.

Number Average Atmospheric Location

Density (cm-3) Diameter (mm) References

10-1-1 1-2 Voigt et al. (2000a, b) Arctic

10-3-10-2 1-4 Dye et al. (1992) a Arctic

10-4-10-2 4-10 Hofmann and Deshler (1991) a Antarctic

10-5-10-3 10-20 Fahey et al. (2001) b; Arctic

Northway et al. (2002a) b

a Note that these measurements are mode diameters of a lognormal distribution. Some particles, most likely containing nitric acid, were observed up to

13 mm in diameter (Wofsy et al., 1990).

b These observations on the average show PSC number densities of about 2 ¥ 10-4 cm-3 with a mode diameter centered near 14 mm, assuming a nitric

acid trihydrate (NAT) composition.
POLAR OZONE

3.24

imate size for the particles based on the sedimentation

speed of the particle layer, rather than directly from the

lidar signal. Their derived particle sizes are in reasonable

agreement with in situ observations (Fahey et al., 2001).

Overall, the assessment of both the Arctic and

Antarctic studies, on observed and inferred PSC particle

sizes, indicates that large nitric acid particles cannot initially nucleate on synoptic-scale ice clouds. However, the

formation mechanism of large nitric acid particles still

remains uncertain. Thus, more laboratory and field studies

are needed to better test which of the above mechanisms

is most likely to dominate the rate of large nitric acid particle production in the polar stratosphere.

Remote

Lidar observations are useful for constraining PSC

particle sizes. Depolarization measurements by lidar can

also provide strong evidence for the presence of solid PSC

particles. As indicated above, only large solid PSC particles

can cause denitrification, and lidar observations provide

valuable information on the horizontal and vertical extents

of solid PSC particle distributions in the stratosphere.

Lidar studies provide new estimates of the occurrence of solid PSC particles, which is an important parameter for constraining microphysical models. Toon et al.

(2000) have reanalyzed lidar observations from all DC-8

flights during the 1989/1990 Arctic winter and find that

large solid PSC particles are more common in the Arctic

stratosphere than the smaller liquid PSC particles (Type

Ib). Three years of lidar observations from Ny Ålesund

(79°N) (Biele et al., 2001) have shown that at least 50%

of PSCs contained solid particles, with some of these

clouds being of Type Ib, normally attributed to pure liquid

clouds. Several studies have also derived new estimates

of solid particle number concentrations. Gobbi et al.

(1998) have analyzed many Antarctic lidar vertical profiles and estimate that solid PSC particles constituted less

than 1% of the available condensation nuclei (therefore,

typically less than 0.1 particles cm-3). Toon et al. (2000)

derived similar values for the 1989/1990 winter Arctic

stratosphere. Biele et al. (2001) and Tsias et al. (1999)

estimate that many depolarizing clouds must typically

contain fewer than about 0.005-0.01 cm-3 solid particles

and that such particles can rarely grow to their equilibrium sizes. In summary, all the lidar studies discussed

here suggest that large solid particles that can cause denitrification are widely distributed in both hemispheres.

Some clouds observed by lidar are consistent with

high number densities of small solid nitric acid particles

(typically >1 cm-3 and <2 mm in diameter) (Tsias et al.,

1999; Toon et al., 2000; Hu et al., 2002). These clouds

are frequently associated with the outflow from mountain

waves (e.g., Hu et al., 2002). Efficient NAT nucleation

on numerous small ice particles formed in wave clouds

could be an important mechanism for generating such a

dense population of small solid PSC particles (Carslaw et

al., 1998a, 1999; Wirth et al., 1999; Larsen et al., 2002).

Some studies suggest that small solid PSC particles generated by wave clouds may also play a central role in

producing large nitric acid particles that lead to denitrification (see the previous section on in situ observations).

Satellite instruments are also capable of observing

PSC particles. Stratospheric Aerosol Measurement

(SAM) II (McCormick et al., 1981), Cryogenic Limb

Array Etalon Spectrometer (CLAES) (Mergenthaler et al.,

1997), Polar Ozone and Aerosol Measurement (POAM)

(Steele et al., 1999; Fromm et al.,1997, 1999; Bevilacqua

et al., 2002), and Improved Limb Atmospheric Sounder

(ILAS) (Kondo et al., 2000; Irie et al., 2001) instruments

have taken many aerosol extinction vertical profiles of

PSCs in the Arctic and Antarctic stratosphere over the last

two decades (see Appendix 3A for satellite data descriptions). Several recent studies have paired up aerosol

extinction measurements with water and nitric acid gasphase measurements onboard the same (Tabazadeh et al.,

2000; Stone et al., 2001; Nedoluha et al., 2000; Dessler et

al., 1999) or different satellites (Santee et al., 2002) to

determine the scale and magnitude of denitrification and

dehydration in both polar regions (see Sections 3.2.2.3

and 3.2.2.4).

3.2.2.2 PARTICLE COMPOSITION

Inferred Particle Compositions

PSCs that exist at temperatures greater than the ice

frost point have long been accepted to be composed of

nitric acid and water, either in the form of a nitric acid

hydrate or as supercooled solution droplets of nitric acid,

sulfuric acid, and water (supercooled ternary solution,

STS). Numerous attempts have been made to infer the

composition of nitric-acid-containing PSCs by comparing

particle volumes (either measured directly or derived from

satellite extinction) and/or gas-phase nitric acid concentrations with equilibrium model calculations assuming

various particle compositions (WMO, 1999; Hopfner et

al., 1998; Steele et al., 1999; Santee et al., 2002; Saitoh et

al., 2002; Strawa et al., 2002). These studies have provided strong support for the existence of STS droplets,

but confirmation of the presence of different nitric acid

hydrates has been more difficult to establish (WMO,

1999). This difficulty may be due to the fact that nitric

acid hydrate particles, which can be considerably larger

than STS droplets, are often not in thermodynamic equilibrium (e.g., Tsias et al., 1999; Biele et al., 2001).
POLAR OZONE

3.25

Direct Determinations of Particle Composition

New in situ observations using an aerosol mass

spectrometer have measured PSC particle composition

(Schreiner et al., 1999, 2002; Voigt et al., 2000a, b; Larsen

et al., 2000, 2002). The mole ratio of nitric acid to water

observed using an aerosol mass spectrometer in a

mountain-induced gravity wave cloud over Scandinavia

(Voigt et al., 2000b) shows close agreement with an STS

droplet composition predicted by a model. The same

aerosol mass spectrometer flown in January 2000 detected

NAT particles, identified by a H2O to HNO3 mole ratio of

3:1 and confirmed to be solid particles from co-located

backscatter measurements (Voigt et al., 2000a; Larsen et

al., 2000); see Figure 3-16. These direct observations of

particle composition in PSCs are important because they

confirm that the thermodynamic models used to predict

liquid aerosol compositions (e.g., Carslaw et al., 1997a)

are reliable, and that NAT, long predicted to exist in the

stratosphere (Hanson and Mauersberger, 1988), actually

does exist there.

3.2.2.3 DENITRIFICATION

Observations of Denitrification

In situ Arctic observations from the ER-2 aircraft

in January to March 2000 detected the most severe and

extensive denitrification observed in the Arctic stratosphere (Popp et al., 2001); see Figure 3-17. Average

removal of as much as 60% of total reactive nitrogen

(NOy) was observed throughout the core of the vortex near

20 km (Figure 3-17). Waibel et al. (1999) have also analyzed balloonborne observations of denitrification in

1995. These observations reveal an approximate 50%

reduction in NOy at 20 km. Using tracer measurements to

eliminate deficits due to mixing (see Section 3.3.1.2),

Waibel et al. (1999) concluded that at least 82% of the

observed NOy loss at 20 km was due to denitrification.

Sugita et al. (1998) and Hintsa et al. (1998) have also

observed similar levels of Arctic denitrification near 20

km, in 1995 and 1996 winters, respectively. Together,

these observations show that some Arctic air parcels were

75000 76000 77000 78000 79000 80000 81000

2 0

2 2

2 4

C Altitude (km)

25 January 2000, Time UT (sec)

2

3

4

5 A 6

STS Condensed phase H O:HNO 2 3 mole ratio

NAT

0.1

1 0

B Particle volume (μm3/cm3

)

Figure 3-16. In situ observations of PSCs using an

aerosol mass spectrometer, showing particles with

a HNO3/H2O ratio consistent with supercooled

HNO3/H2SO4/H2O droplets. Panel A displays the

measured molar ratios, panel B shows particle

volume, and panel C shows the balloon altitude. The

red color indicates the presence of nitric acid trihydrates (NAT), blue indicates supercooled ternary

solution (STS) particles, and gray represents particles with insufficient information to derive molar

ratios. Adapted from Voigt et al. (2000a).

20

18

16

14

12

10 Altitude (km)

15 10 5 0 -5 -10 -15

Denitrification (ppbv) and NOy

* (ppbv)

NOy

*

Figure 3-17. Vertical profiles of redistributed NOy in

the 1999-2000 Arctic vortex. Positive values of

denitrification represent NOy removal. Blue and red

symbols represent measurements made at greater

than 70°N and less 70°N equivalent latitude, respectively. The solid black line representing NOy* (the

maximum available NOy at a given altitude) was

determined from altitude-bin-averaged N2O values

for measurements made throughout the winter at

equivalent latitudes greater than 70°N. Maximum

ER-2 flight altitudes of 20 to 12 km represent a

potential temperature of approximately 460 to 475

K. Adapted from Popp et al. (2001).
severely denitrified (>50%) near 20 km in several cold

winters.

Figure 3-18 shows the time evolution of HNO3 in

both hemispheres for a number of winters on three potential temperature surfaces, based on Microwave Limb

Sounder (MLS) observations (see Appendix 3A for MLS

data description). Large irreversible depletions (>80%)

in HNO3 vapor amounts occur only in the SH (Santee et

al., 1999), but much smaller irreversible depletions

(~20%) have also been observed in cold Arctic air parcels

(Dessler et al., 1999; Santee et al., 2000, 2002). However,

large permanent depletions in HNO3 concentrations

(>40%) have been observed in cold Arctic winters by

satellite instruments with finer vertical resolution than

MLS (~6-km vertical resolution), such as the ILAS instrument, which has a 1- to 2-km vertical resolution (Kondo

et al., 2000; Irie et al., 2002; see also Appendix 3A for

satellite data descriptions). Thus, extensive denitrification does occur in both hemispheres based on both in situ

and remote sensing observations. However, severe denitrification in the Arctic must be narrow in depth because

if the denitrified layers were deeper than 6 km, then MLS

would have been able to detect this irreversible loss in

gas-phase nitric acid (Tabazadeh et al., 2001). Individual

nitric acid vertical profiles from the ILAS instrument

(Kondo et al., 2000; Irie et al., 2002) during the winter of

1996/1997 further show that the vertical range of denitrification in the Arctic is typically ~2 to 3 km deep (from

about 18 to 21 km), with immediate nitrification occurring below this altitude range.

Tabazadeh et al. (2000) have shown that denitrification over a broad altitude range (>10 km in depth)

occurs rapidly in the Antarctic when the duration of an

average PSC event is about 2 weeks (defined to be the

time spent by an isentropic air parcel below 195 K at 450-

K potential temperature). Tabazadeh et al. (2000) suggest

that deep extensive denitrification currently occurs only

in the Antarctic because average Arctic PSC events last

nearly half as long as those in the Antarctic. Antarctic

denitrification rapidly occurs during the mid- to late-June

time period (see Figure 3-18). The short persistence of

PSC events in the NH limits the extent of severe denitrification to only a few kilometers, as observed in many past

cold Arctic winters (Popp et al., 2001; Sugita et al., 1998;

Hintsa et al., 1998; Kondo et al., 2000; Irie et al., 2001,

2002). In Section 3.3.4, the effect of denitrification depth

on ozone loss is assessed for both hemispheres.

Model Simulations of Denitrification

Improvements in our understanding of PSC particle sizes and number concentrations (Section 3.2.2.1)

now allow for more sophisticated treatments of denitrification in models. Current 3-D chemical transport models

include highly simplified representations of denitrification that are now recognized to be incorrect. These models

assume that nitric acid is carried downward on sedimenting ice particles wherever temperatures are lower

than the ice frost point (Chipperfield et al., 1993;

Considine et al., 2000). Removal of nitric acid on sedimenting ice particles is clearly not operating in the Arctic

because models, using this assumption, have been unable

to produce any denitrification by this mechanism

(Chipperfield and Pyle; 1998; Davies et al., 2002).

Drdla et al. (2002) have used a coupled

microphysical-chemistry trajectory model to show that

synoptic-scale ice clouds could not have caused the massive denitrification observed during the 1999/2000 winter.

They concluded that the large nitric acid particles that

caused denitrification were not nucleated on ice particles.

Jensen et al. (2002) have used a one-dimensional

version of the Community Aerosol and Radiation Model

for Atmospheres (CARMA) microphysics cloud model to

show that solid nitric acid number densities in the range

of 10-2 to 10-3 cm-3 are the most efficient in causing rapid

denitrification. The CARMA cloud model has also been

used to show (see Figure 4 in Tabazadeh et al., 2001) that

the vertical range over which denitrification occurs is normally quite deep in the Antarctic (>10 km) but limited in

the Arctic, in general agreement with many observations

of denitrification in both hemispheres (i.e., Santee et al.,

1999; Popp et al., 2001; Sugita et al., 1998; Hintsa et al.,

1998; Kondo et al., 2000; Irie et al., 2001, 2002). The

Arctic denitrified layers are shallow in depth mainly

because the stratosphere, even in cold winters, is too warm

below about 17 km for nitric acid hydrate particles to exist.

Waibel et al. (1999) simulated denitrification by

assuming that ice acted as the nucleus for NAT formation.

Most of the denitrification calculated in the model was

caused primarily by sedimentation of NAT particles, with

a fixed number concentration of 5 ¥ 10-3 cm-3, released

upon ice evaporation. In these 3-D model simulations, in

which a simplified treatment of horizontal transport was

used, NAT particles were assumed to be in equilibrium

with gas-phase nitric acid.

Davies et al. (2002) have included a parameterization in a 3-D CTM of NAT particle sedimentation

assuming sizes and number concentrations representative

of those observed by Fahey et al. (2001). NAT particles

were assumed to be in equilibrium with gas-phase nitric

acid in these simulations. The calculated denitrification

magnitude was similar to that observed in winter

1999/2000. In contrast, the modeled denitrification was

significantly less than observed when NAT was allowed

POLAR OZONE

3.26
POLAR OZONE

3.27

 0

 4

 8

12

16 520 K

Arctic MLS HNO

 0

 4

 8

12

16 (ppbv)

465 K

 0

 4

 8

12

16 420 K

1 Oct 1 Jan 1 Apr 1 Jul 1 Oct

 0.0

 0.5

 1.0

 1.5

 2.0 520 K

Arctic MLS ClO

 0.0

 0.5

 1.0

 1.5

 2.0 (ppbv)

465 K

 0.0

 0.5

 1.0

 1.5

 2.0 420 K

1 Oct 1 Jan 1 Apr 1 Jul 1 Oct

 0

 4

 8

12

16 520 K

Antarctic MLS HNO

 0

 4

 8

12

16 (ppbv)

465 K

 0

 4

 8

12

16 420 K

1 Apr 1 Jul 1 Oct 1 Jan 1 Apr

 0.0

 0.5

 1.0

 1.5

 2.0 520 K

Antarctic MLS ClO

 0.0

 0.5

 1.0

 1.5

 2.0 (ppbv)

465 K

 0.0

 0.5

 1.0

 1.5

 2.0 420 K

1 Apr 1 Jul 1 Oct 1 Jan 1 Ap

1991-92 1992-93 1993-94 1994-95 1995-96 1996-97 1999-2000

Figure 3-18. Time series of Upper Atmosphere Research Satellite (UARS) MLS HNO3 (left, ppbv) and ClO (right,

ppbv) at 520 K (~21 km), 465 K (~19 km), and 420 K (~17 km) for both the Northern (top) and Southern (bottom)

Hemispheres. Values shown are averages of data points with solar zenith angle <88° (daylight) in the 70°-80°

equivalent latitude bin. These values represent the best estimates of the abundances at these particular potential temperature surfaces based on the MLS measurements; it should be borne in mind, however, that correlations between these levels may have been introduced either during the retrieval process (since the retrieval

pressure grid is finer than the width of the averaging kernels) or through interpolation of the retrieved data onto

potential temperature surfaces, and thus the true number of completely independent pieces of information is likely to be less than the number of surfaces shown here. For the daily averages in this equivalent latitude bin the

precision is roughly 0.15 ppbv for HNO3 and 0.05 ppbv for ClO. Different years are represented by different colors as indicated in the legend. The x-axis tick mark increment is 10 days; dotted vertical lines demark calendar

months. The 1991 to 1997 and 2000 data are adapted from Santee et al. (1999) and (2000), respectively.
POLAR OZONE

3.28

to nucleate only on ice particles. This result is consistent

with previous simulations (Carslaw et al., 2002; Drdla et

al., 2002) showing that the observed large nitric acid particles could not have nucleated on synoptic-scale ice

clouds during the winter of 1999/2000.

Mann et al. (2002) have developed a 3-D model of

denitrification that takes into account the time dependence of both growth and sedimentation of large NAT

particles. In these simulations, NAT particle growth times

are on the order of several days, and therefore the magnitude of denitrification strongly depends on individual

particle growth cycles. In fact, the extent of modeled denitrification is strongly amplified when areas of low temperature are stable and concentric with the vortex,

allowing for individual particles to persist for a longer

time and to grow to larger sizes. Thus the area of low temperatures by itself is perhaps not the best indicator for

predicting the severity of denitrification in the Arctic.

In summary, new model calculations show that the

long-accepted mechanism of nitric acid removal on sedimenting ice particles cannot account for observed levels

of denitrification. Denitrification is most likely caused

by sedimentation of large nitric acid particles. Various

models are able to reproduce, in broad terms, the observed

levels of denitrification by assuming sedimentation of

large nitric acid particles. In addition, recent model calculations show that Arctic denitrified layers are normally

shallow in depth because the stratosphere is too warm

below about 17 km to allow for large particles to carry

much nitric acid beyond this altitude. Establishing the

formation mechanism of the large nitric acid particles will

be important for the development of realistic models of

denitrification. The sensitivity of ozone loss to denitrification is discussed in Section 3.3.4.

3.2.2.4 DEHYDRATION

Figure 3-19 shows the time evolution of water

vapor in the SH during 1998 (Nedoluha et al., 2000).

Severe dehydration is observed over a 10-km altitude

range. Similar results are also obtained from analysis of

MLS water vapor data during the Antarctic winter of 1992

(Stone et al., 2001; see also Appendix 3A for MLS data

description). In the Arctic, removal of ~1 ppm of water

vapor over a 1- to 2-km altitude range has been observed

by the ILAS instrument (Pan et al., 2002). The lack of

extensive and deep dehydration in the Arctic is also supported by in situ observations (Vömel et al., 1997; Hintsa

et al., 1998; Herman et al., 2002). Overall, the Arctic

climate, even in cold years, is too warm to allow for formation of widespread persistent ice clouds that lead to

dehydration. The sensitivity of ozone loss to dehydration

is discussed in Section 3.3.4.

3.2.3 Polar Ozone Chemistry

The chemical loss of polar ozone during winter and

spring occurs primarily by two gas-phase catalytic cycles

that involve halogen oxide radicals:

Cycle 1

(1a) ClO + ClO + M Æ ClOOCl + M

(1b) ClOOCl + hn Æ 2 Cl + O2

(1c) 2 (Cl + O3 Æ ClO + O2)

Net: 2 O3 Æ 3 O2 (3-1)

Cycle 2

(2a) BrO + ClO Æ Br + Cl + O2

(2a¢) Æ BrCl + O2

(2b) BrCl + hn Æ Br + Cl

(2c) Br + O3 Æ BrO + O2

(1c) Cl + O3 Æ ClO + O2

Net: 2 O3 Æ 3 O2 (3-2)

M J J A S O N D

1998

10

20

30

40

50 Altitude (km)

2

4

6

8

ppmv

descent

mid-latitude air

dehydration

Figure 3-19. The water vapor mixing ratio as measured by POAM in the Southern Hemisphere from 27

April 1998 to 12 December 1998. Profiles are

obtained from daily averages of the measurements

from POAM, which generally provides 14 measurements per day. Measurements from both inside and

outside of the vortex are included. The POAM measurement latitude biannually cycles from 88°S at the

equinoxes (March 21 and September 21) to 62°S at

the solstices (June 21 and December 21). Adapted

from Nedoluha et al. (2000).
POLAR OZONE

3.29

Abundances of chlorine monoxide (ClO) in the

polar vortex are greatly elevated by reactions of inactive

chlorine reservoir species on various types of polar stratospheric clouds (PSCs) that form when temperatures drop

below about 195 K (see Section 3.2.2). Abundances of

bromine monoxide (BrO) determine the removal rate by

Cycle 2, which contributes about 50% to the total chemical loss rate of polar ozone (Chipperfield and Pyle, 1998).

In contrast to ClO, the abundance of BrO is not strongly

affected by reactions involving PSCs because less than

half of the available inorganic bromine budget is

sequestered in reservoirs such as bromine nitrate (BrNO3)

and hydrogen bromide (HBr) prior to processing.

Since the previous Assessment, new studies have

addressed the rates of ozone destruction by Cycles 1 and

2 in an effort to reconcile apparent discrepancies between

measured and modeled chemical loss rates of Arctic ozone

(see Section 3.3). The chlorine and bromine cycles are

discussed separately below in Sections 3.2.3.1 and 3.2.3.2.

Model calculations reveal that, at present and for the foreseeable future, winter polar ozone loss will be dominated

by reactions involving ClO and BrO (e.g., Chipperfield

and Pyle, 1998; Shindell et al., 1998b).

3.2.3.1 CHLORINE

A number of modeling studies have had difficulty

accounting for observed chemical loss rates of Arctic

ozone, particularly during midwinter when insolation is

weakest (see Section 3.3). As a consequence, attention

has focused on reducing uncertainties in key rate parameters, in particular those for formation and photolysis of

the chlorine oxide dimer (ClOOCl), i.e., parameters k1a

(first-order reaction-rate constant) and J1b (photolysis

rate). These reactions determine the rate of ozone loss by

Cycle 1, such that uncertainties in these kinetic parameters couple directly into the uncertainties in modeled

ozone loss rates. Results of these studies are summarized

in the section below on chlorine photochemistry.

The loss of polar ozone via Cycles 1 and 2 is determined by the temporal evolution of ClO. Numerous

studies have examined the seasonal variations of inorganic chlorine partitioning, including production of

high abundances of reactive chlorine (ClOx, defined as

[ClO] +2 [ClOOCl]) by PSCs (“activation”); maintenance

of high abundances of ClOx throughout the winter; and

deactivation of ClOx back to hydrogen chloride (HCl) and

chlorine nitrate (ClNO3) (“recovery”). These results are

summarized in the section below on chlorine seasonal

evolution.

Simultaneous measurements of most of the major

inorganic and organic chlorine species provide the ability

to examine the overall chlorine budget for the Arctic and

Antarctic stratospheres during periods of rapid ozone loss.

These investigations can assess the possible role of species

not considered in standard models, such as higher oxides

of chlorine. The results of these studies are summarized

below (the section on the chlorine budget).

Chlorine Photochemistry

Our understanding of the key kinetic parameters

governing ozone loss by Cycle 1 (k1a and J1b) has improved through new laboratory studies and atmospheric

observations. Bloss et al. (2001) found k1a to be up to 25%

larger than the value recommended in the Jet Propulsion

Laboratory (JPL) Publication 00-3 compendium (Sander

et al., 2000; hereafter referred to as JPL 00-3 in this

chapter) at temperatures below 210 K. The impact of this

increase on calculated ozone loss rates depends on the

manner in which the models treat ClO. In the case of

models constrained by observed abundances of ClO, the

rate of ozone loss due to Cycle 1 is roughly proportional

to the increase in k1a. For models that allow ClO and

ClOOCl to repartition within the constraint of constant

active chlorine, the total loss rate is largely independent

of this change in k1a, because ClO abundances respond in

the opposite sense to the change in the rate constant

(Figure 3-20). Thus, the maximum effect of an increase

in k1a based on the results of Bloss et al. (2001) is to

increase total ozone loss rates by no more than 25% under

cold polar conditions. Such an increase is insufficient to

resolve fully the factor-of-2 discrepancies between modeled and measured Arctic ozone loss rates during January

found in several recent studies (e.g., Woyke et al., 1999;

Becker et al., 2000).

The recommended absorption cross section for

ClOOCl has not changed since the previous Assessment.

However, there have been several important laboratory

and theoretical studies (Moore et al., 1999; Kaledin and

Morokuma, 2000; Toniolo et al., 2000) of the product

yields from photolysis of ClOOCl at wavelengths longer

than 300 nm, the spectral region that contributes most to

the overall photolysis rate of ClOOCl. These new results

reduce a significant uncertainty in our knowledge of ozone

loss rates. Only the production of chlorine atoms and

chloroperoxy radical (ClOO), leading to 2Cl + O2 upon

the rapid thermal decomposition of ClOO at polar temperatures, results in catalytic loss of ozone; production of

ClO + ClO from ClOOCl photolysis leads to a null cycle

that has no effect on ozone. Moore et al. (1999) recently

reported that two chlorine atoms and molecular oxygen

(O2) are the primary products of photolysis of ClOOCl at

248 nm and at 308 nm. The ClOO product rapidly decomposes to atomic chlorine (Cl) and O2 due to excess vibrational energy. This conclusion is further supported by
POLAR OZONE

3.30

electronic structure calculations (Kaledin and Morokuma,

2000; Toniolo et al., 2000). These studies reduce the

uncertainties of the product yield of ClOOCl photolysis

at key wavelengths for chemical loss of polar ozone relative to earlier laboratory investigations.

Several new sets of atmospheric observations provide quantitative tests of our understanding of k1a and J1b.

Recent balloonborne in situ observations of the rate of

decay of ClO immediately after sunset in the Arctic vortex

(Vömel et al., 2001) are explained better by the larger

value for k1a at temperatures near 190 K than by the value

recommended by JPL 00-3 (Figure 3-21). An analysis

of data obtained during five winters (1996 to 2000)

(Solomon et al., 2002) found good agreement between

measurements of column ClO and model calculations

employing the value of J1b / k1a from JPL 00-3, in contrast

to an earlier report, using only data for 1996, that indicated the ratio J1b / k1a might be 50% too low (Solomon et

al., 2000). A similar conclusion had been reached in a

study of earlier ground-based ClO observations over

Antarctica (Shindell and de Zafra, 1995).

The chemical loss rate of ozone, particularly during

early winter, is also sensitive to the dependence of J1b on

solar zenith angle (SZA). Avallone and Toohey (2001) concluded, based on an examination of in situ observations

of ClO obtained at SZAs between 79° and 90°, that the

photolysis rate of ClOOCl varied with SZA in a manner

generally consistent with recommended absorption cross

sections for ClOOCl. An analysis of the data they presented using the larger Bloss et al. (2001) value for k1a

yields values for J1b that agree well with photolysis rates

using JPL 00-3 cross sections. The significantly smaller

value for J1b based on absorption cross sections of Huder

and DeMore (1995) appears to be inconsistent with these

results and those of Shindell and de Zafra (1995), Raffalski

et al. (1998), Stachnik et al. (1999), and Solomon et al.

(2002). Furthermore, there is no evidence from any of

these studies that supports the notion that ClOOCl may

photolyze at an appreciable rate in optically thin spectral

regions (i.e., >420 nm). Such a process could enhance

ozone loss rates at high SZAs and account for some of the

discrepancy between measured and modeled ozone loss

rates in midwinter (Rex et al., 2002b; see also Section 3.3).

Chlorine Seasonal Evolution

Another element critical in accounting for ozone

loss is the temporal evolution of the ozone-destroying

halogen radicals. Mixing ratios of ClO remain elevated

(~1 to 2 parts per billion by volume (ppbv)) from

May/June until September (Figure 3-18) for all years over

Antarctica for which observations are available (WMO,

1995; WMO, 1999; Santee et al., 2000; Wagner et al.,

2001, 2002; Solomon et al., 2002), in the region where

temperatures between 14 and 24 km remain very low

Figure 3-20. (a) Calculations of 24-hour average

total chemical ozone loss rate for an air mass at

70°S, 50 hPa, as a function of temperature for midSeptember conditions (e.g., ClO + 2◊ClOOCl = 2

ppb; BrO + BrCl = 10 ppt; O3 = 2 ppm; declination =

5°S) assuming the following: case 1, JPL 00-3

(Sander et al., 2000) kinetics (solid line); case 2, the

Bloss et al. (2001) value for k1a and the noontime

concentration of ClO fixed at the same value used

in case 1 (dotted line); and case 3, the Bloss et al.

value for k1a and the noontime concentration of ClO+

2◊ClOOCl fixed at the same value used in case 1

(dashed line). (b) Same as (a), except chemical loss

rates due to just the ClO + ClO cycle are shown.

Case 2 approximates the effect of the Bloss et al.

value for k1a on a model calculation constrained by

measured ClO. Case 3 approximates the effect of

Bloss et al. k1a within a free-running CTM or GCM,

where ClO and ClOOCl are allowed to equilibrate.

Adapted from Bloss et al. (2001).
POLAR OZONE

3.31

(below PSC thresholds) for several months (Figure 3-13,

bottom). In addition, the southern polar vortex remains

intact well into the spring season. Under these conditions,

the total amount of ozone destroyed over Antarctica is

nearly complete for an 8- to 10-km-thick altitude layer, a

condition that is relatively insensitive to the chemical loss

rate at contemporary abundances of inorganic chlorine

(WMO, 1999).

Three-dimensional CTMs are able to simulate the

seasonal evolution of ClO in the Antarctic polar vortex

remarkably well (Ricaud et al., 1998; Solomon et al.,

2000, 2002). Figure 3-22 compares ground-based column

measurements of ClO above Scott Base, Antarctica

(77.8°S), with calculations from the SLIMCAT model.

Comparisons for the ClO mixing ratio at 480 K are also

shown. During mid- to late winter, the rise of ClO is determined primarily by increasing solar illumination at

midday. The good agreement between theory and observations of column ClO during this time period suggests

that the altitude range over which chlorine is activated is

well reproduced by the model. During early spring, the

short-term fluctuations in ClO are related to movement of

the vortex over Scott Base, and the longer-term decline in

ClO is the result of recovery into the reservoirs HCl (the

primary sink for ClOx in the denitrified Antarctic vortex)

Figure 3-21. Derived rate of decay of ClO (points

with 1s error bars) from atmospheric observations

of ClO obtained in the Arctic vortex during SOLVETHESEO 2000 compared with the product of M◊k1a

from the JPL 00-3 (Sander et al., 2000) recommendation (thin solid line) and from the laboratory study

of Bloss et al. (2001) (thick solid line). Uncertainties

for the JPL 00-3 rate are given by the dashed lines.

Uncertainties for the laboratory measurement of

Bloss et al. (2001) are not shown for clarity, but they

are approximately the same size as the error bars

for the results based on the ClO observations.

Adapted from Vömel et al. (2001).

228 235 243 252 263 274

Day Number

Aug 16 23 31 Sep 9 20 Oct 1

0

0.5

1.0

1.5

2.0

2.5

3.0 ClO Column Density (x 1015 cm-2)

Retrieval

Model - 2000

228 235 243 252 263 274

Day Number

0.0

0.5

1.0

1.5

2.0

2.5 Mixing Ratio (ppbv)

Aug 16 23 31 Sep 9 20 Oct 1

Model - 2000

Measurements

Potential Temperature: 480 K

Altitude ~ 21 km

Figure 3-22. Top panel: Retrieved ClO column density (day-night difference) for 11 August through 4

October 2000 over Scott Base, Antarctica (78°S)

(blue line), and the same quantity calculated using

the SLIMCAT model assuming JPL 97-4 (DeMore et

al., 1997) kinetics (red line). Bottom panel: Same

as top panel, except for ClO mixing ratio (day-night

difference) at 480 K. Adapted from Solomon et al.

(2002).
POLAR OZONE

3.32

and ClNO3 (Solomon et al., 2002). As a result of widespread suppression of gas-phase HNO3 in SLIMCAT, high

abundances of ClO are sustained throughout the early

spring season (September) until ozone is nearly completely removed. In October, HCl is observed to reappear

much faster than ClNO3 in the core of the vortex due to

the shift in partitioning of Cl/ClO and nitric oxide/nitrogen

dioxide (NO/NO2) to favor Cl and NO driven by exceedingly low ozone, while ClO recovers mainly to ClNO3 in

the edge region of the Antarctic vortex (e.g., Douglass et

al., 1995; Ricaud et al., 1998). Because most models simulate these features reasonably well, they are able to

account for Antarctic ozone loss in a quantitative manner.

The situation for the Arctic winter is quite different,

because chemical ozone loss depends more critically upon

the details of chlorine activation and deactivation, and the

timing of the breakup of the northern polar vortex. During

cold Arctic winters, high levels of ClO are observed

throughout the polar vortex (Raffalski et al., 1998;

Stachnik et al., 1999; Klein et al., 2000; Santee et al.,

2000). Considerably more year-to-year variability is seen

in Arctic measurements of ClO compared with Antarctic

data, and peak values of Arctic ClO for cold winters are

somewhat lower than observed in the Antarctic (Figure 3-

18). Recent GOME measurements of chlorine dioxide

(OClO) (Figure 3-23), which indicate much greater yearto-year variability in active chlorine for the Arctic as well

as considerably higher levels of active chlorine for the

Antarctic (Wagner et al., 2001, 2002), provide a picture

consistent with the MLS observations of ClO. The GOME

observations of Antarctic OClO are also consistent with

earlier ground-based observations of OClO (Miller et al.,

1999; see also Appendix 3Afor satellite data descriptions).

The MLS and GOME measurements show that

elevated levels of ClOx in the Arctic, even for cold years,

decline rapidly in early spring, in contrast to the Antarctic,

where high ClOx persists well into spring (Figures 3-18 and

3-23). Consequently, the total quantity of ozone destroyed

0

2

3 Antarctic

winters

1 9 9 5/9 6

1 9 9 6/9 7

1 9 9 7/9 8

1 9 9 8/9 9

1 9 9 9/0 0

2 0 0 0/0 1 OClO SLANT COLUMN DENSITY (SCD) 1014 molecules/cm2

DATE

Arctic

Winters

1-Nov 1-Dec 31-Dec 30-Jan 1-Mar 31-Mar 30-Apr

(Antarctic values are shifted by 6 months)

Figure 3-23. Daily maximum slant column density (SCD) of OClO, at SZA of 90°, measured by GOME for a

number of Arctic and Antarctic winters. Data for individual Arctic winters are shown, as indicated, whereas the

range of data for Antarctic winters is shown by the shaded region. The time associated with the Antarctic

measurements has been shifted by 6 months to facilitate comparison with Arctic observations at comparable

seasons. The air mass factor (ratio of SCD to vertical column density) depends on the height profile of OClO,

which is not known a priori. Consequently, the SCD is the preferred choice for comparison of Arctic and

Antarctic data. Adapted from Wagner et al. (2001, 2002).
POLAR OZONE

3.33

in the Arctic vortex depends strongly on the rate of chlorine deactivation, which in turn is related to the extent of

denitrification (e.g., Rex et al., 1997; Waibel et al., 1999;

Tabazadeh et al., 2000) and the efficiency of chlorine reactivation (e.g., Solomon, 1999; Hanisco et al., 2002; Drdla

and Schoeberl, 2002). Abundances of ClOx over the Arctic

decrease rapidly when temperatures increase above

~200 K, because of photochemical release of NOx from

nitric acid that remains in excess of reactive chlorine

throughout the winter. During this recovery period, observations have shown that ClNO3 is the primary inorganic

chlorine species, representing >80% of the available chlorine (Chapter 3 of WMO, 1995). It has long been assumed

that chlorine can be readily reactivated on PSCs during this

recovery period (provided temperature drops below ~195

to 200 K), leading to significant additional ozone loss.

However, a recent study based on analyses of in situ observations of hydroxyl radical (OH) and hydroperoxyl radical

(HO2) (Hanisco et al., 2002) concludes that key heterogeneous reactions that reactivate chlorine proceed more

slowly than currently recommended rates. The consequences of this finding have yet to be explored in photochemical model studies of Arctic ozone loss.

Most important, however, 3-D chemistry and transport models (Ricaud et al., 1998; Massie et al., 2000; van

den Broek et al., 2000) as well as trajectory simulations

(Woyke et al., 1999; Danilin et al., 2000) are able to simulate well the high levels of ClO observed in the Arctic,

indicating that chlorine activation schemes used in

photochemical models are relatively accurate in describing

large-scale features of chlorine activation. Recent model

simulations also suggest that the rate and extent of halogen

activation in the polar vortex are not as sensitive to PSC

composition as previously thought (Carslaw et al., 1997b;

Becker et al., 1998; Woyke et al., 1999; Danilin et al., 2000).

This lack of sensitivity arises because most heterogeneous

halogen activation rates are much faster at low temperatures than deactivation rates of ClOx for air parcels outside

of PSCs (e.g., Solomon, 1999).

Although models simulate the seasonal evolution of

ClO in the Arctic reasonably well for cold winters, they

have some difficulty for warm winters, where minimum

temperatures are close to the threshold for formation of

PSCs (e.g., Klein et al., 2000). For the winter of 1998/1999,

ground-based observations of ClO from Ny Ålesund,

Spitzbergen (78.9°N), revealed little or no enhancements

above background levels, whereas the SLIMCAT model

predicted ClO mixing ratios as high as 1.0 ppbv. This discrepancy has been attributed to a small cold bias (~1 K) in

the UKMO temperatures input to the SLIMCAT model

(Klein et al., 2000; Knudsen et al., 2002), although it is also

possible that the PSC nucleation scheme in SLIMCAT is

unrealistic at temperatures near NAT thresholds. In either

case, such a problem highlights the extraordinary sensitivity of Arctic ClO to temperatures and microphysics

schemes for winters where the minimum temperatures are

very close to the threshold for formation of PSCs.

Chlorine Budget

Ideally, an assessment of the chlorine budget

should be based on simultaneous measurements of the

primary inorganic (e.g., HCl, ClNO3, ClO, ClOOCl) and

organic (e.g., chlorofluorocarbons and other chlorinecontaining halocarbons) chlorine species. A significant

number of studies, all of which lack observations of

ClOOCl, have indicated good agreement between the

inorganic and organic chlorine budget for the lower polar

stratosphere (e.g., von Clarmann et al., 1995; Engel et al.,

1997; Mickley et al., 1997; Ricaud et al., 1998; Michelsen

et al., 1999; Pierson et al., 1999). For example, balloonborne microwave and whole-air sampler measurements

of [HCl] + [ClO] + 2[ClOOCl*] versus N2O for air in the

core of the Arctic vortex on 27 January 1995 (Stachnik et

al., 1999), made under conditions of highly elevated ClO,

agree well with estimates of inorganic chlorine based on

the measured decomposition of organic halocarbons. The

above-mentioned studies relied on calculated concentrations of ClOOCl, termed ClOOCl*, assuming a steadystate relation with measured ClO. Their validity is

supported by the good agreement between laboratory rates

for k1a and J1b (needed to calculate ClOOCl*) and values

inferred for these kinetic parameters from the time rate of

change of ClO in the atmosphere (Avallone and Toohey,

2001; Vömel et al., 2001).

Also, these findings are consistent with reasonably

good agreement (differences of about ±15%) between the

disappearance of organic chlorine (CCly) species and the

appearance of inorganic chlorine (Cly) species observed

for the summer polar stratosphere (Section 3.2.4.3), a

region of the atmosphere for which the contribution to

Cly is dominated by HCl and ClNO3.

3.2.3.2 BROMINE

Cycle 2 (BrO + ClO; see beginning of Section

3.2.3) makes important contributions to polar ozone loss.

As shown below in the section on bromine monoxide

abundances, there is now reasonably good agreement

between measurements of BrO obtained by various techniques. This is a significant advance in our understanding

because important differences had been noted in the previous Assessment (WMO, 1999). This convergence of

measurements allows for fairly accurate assessment of the

contribution of bromine to chemical loss of polar ozone.
POLAR OZONE

3.34

Profiles of inorganic bromine (Bry) based on measurements of BrO have recently been compared with estimates based on the observed falloff (with increasing

height) of the organic source species. These comparisons,

discussed in the section on bromine trends and budget,

show a slight offset that may result from either direct influx

of ~3 parts per trillion by volume (pptv) of inorganic

bromine across the tropical tropopause or some organic

species not accounted for. Finally, profile measurements

of BrO discussed also in the bromine budget section have

been used to determine trends in total bromine loading that

can be compared with trends based on the organic bromine

(CBry) content of the lower atmosphere. 

Bromine Monoxide (BrO) Abundances

Harder et al. (1998) compared in situ BrO measurements from the ER-2 aircraft and a balloon flight

with profiles of BrO obtained by the Differential Optical

Absorption Spectroscopy (DOAS) technique (Figure 3-

24). There is a systematic difference in these two sets of

observations, where DOAS measurements are somewhat

larger than in situ, although this is within the combined

uncertainties of the measurements. Consequently, the estimates of the inorganic bromine budget based on these sets

of measurements have ranged from ~16 pptv (in situ) to

~20 pptv (DOAS) (Avallone et al., 1995; Pfeilsticker et

0 5 10 15 20

0

5

1 0

1 5

2 0

2 5

3 0

3 5

4 0

Arctic/Antarctic: 

McKinney et al. 1997 

 Mid-Latitudes: 

 

SLIMCAT Model: Leon

Kiruna

Harder et al. 1998: Leon, Nov 1996

Kiruna, Feb 1997

Gap, July 1997

Brune et al. 1988

Toohey et al. 1990

Avallone et al. 1995

Brune et al. 1989

Toohey et al. 1990

Avallone et al. 1995

Pundt 1997

MIXING RATIO (ppt) ALTITUDE (km) Total Br Schauffler et al. 1998 Total Br Wamsley et al. 1998

Figure 3-24. Measurements of profiles of BrO obtained at midlatitudes, in the Arctic and the Antarctic, by the

in situ resonance fluorescence technique (red and blue open symbols and blue * symbol) and by the remote

Differential Optical Absorption Spectroscopy (DOAS) technique (closed symbols). Also shown are two estimates of total inorganic bromine (total Bry) loading at the top of the atmosphere based on measured abundances of brominated source gases (thin vertical lines). Calculations of profiles for BrO from the SLIMCAT

model, appropriate for conditions of the DOAS observations of BrO from León, Spain (42.6°N), and Kiruna,

Sweden (67.9°N), are shown by the red solid and blue dashed lines. See Harder et al. (1998) for the full citation of the references given on the figure. Adapted from Harder et al. (1998).
POLAR OZONE

3.35

al., 2000). It is important to note that, because these sets

of observations were obtained 5 years apart, more than

half of this difference can be explained by trends in the

bromine source gases, as discussed in the next section.

Sinnhuber et al. (2002) compared ground-based

zenith sky measurements obtained at 11 sites with simulations from the SLIMCAT model in an effort to examine

the detailed processes that govern the partitioning of BrO.

Comparisons for three sites are shown in Figure 3-25. The

simulated abundances of BrO generally agree to within

~10% of the observations over a wide range of seasons,

latitudes, and solar zenith angles. The results are consistent with a total stratospheric bromine loading (sum of

organic and inorganic) of 20 ± 4 pptv, in agreement with

the values deduced from previous remote measurements

of BrO.

The SLIMCAT model tends to overestimate BrO

column abundances at high latitudes, typically when ClO

abundances are elevated (Figure 3-25). Conversely, Friess

et al. (1999) find a discrepancy in the opposite sense

between BrO slant column measurements made at Kiruna,

Sweden (67.9°N), in winter and SLIMCAT model calculations that use JPL 97-4 kinetics (DeMore et al., 1997;

hereafter referred to as JPL 97-4 in this chapter) and a

bromine loading of 20 pptv (the model underestimates

midday measured BrO columns by 20 to 40%). Sinnhuber

et al. (2002) note that the discrepancy highlighted in their

study can be reduced by increasing the rate constant for

reaction 3-2a¢ to the upper limit of the uncertainty of the

JPL 00-3 recommendation. Friess et al. (1999), however,

report that the discrepancy they found is evidence for several pptv of BrO in the free troposphere. The Friess et al.

(1999) interpretation is consistent with interpretations

based on other remote observations of BrO (e.g., Harder

et al., 1998; Fitzenberger et al., 2000).

The discrepancies outlined above have a relatively

minor impact on ozone loss rates calculated directly from

Cycle 2 or on observed abundances of ClO and BrO in the

polar vortices. However, they do raise questions about

the completeness of our understanding of coupled

bromine/chlorine chemistry. Similar questions have been

raised based on aircraft observations of BrO at 20 km (in

situ) and 12 km (remote) within the perturbed polar vortex.

Specifically, Avallone and Toohey (2001) report that

mixing ratios of BrO did not drop to near-zero as expected

with increasing SZA after sunset, when reservoir species

like bromine nitrate (BrNO3), bromine chloride (BrCl),

and hypobromous acid (HOBr) are expected to sequester

nearly all available reactive bromine. Similarly, Wahner

and Schiller (1992) previously reported non-zero BrO

column abundances above 12 km in darkness that were

difficult to explain. Avallone and Toohey (2001) suggest

that thermal decomposition of a weakly bound molecule,

such as BrOOCl (bromochloroperoxide), may be able to

maintain a few pptv of BrO following sunset, but note that

such a process would have little impact on ozone loss rates

because of the rapid decline of ClO at sunset. The existence of adducts of bromine and chlorine oxides has

been postulated in theoretical studies (Gleghorn, 1997;

Bridgeman and Rothery, 1999; Gomez and Pacios, 1999;

Papayannis et al., 2001) and has been observed in an argon

matrix (Johnsson et al., 1995).

The previous Assessment noted the spectroscopic

detection of bromine dioxide (OBrO) in the midlatitude

stratosphere, with implied mixing ratios as high as 20 pptv

(Renard et al., 1997). As such, OBrO would be the dominant nighttime reservoir for inorganic bromine in the midlatitude stratosphere. The same group has since reported

the presence of smaller amounts of OBrO in the nighttime, polar stratosphere (Renard et al., 1998). However,

abundances of even a few hundredths of a part per trillion

(ppt) of OBrO in the nighttime stratosphere are contrary

to our present understanding of bromine photochemistry

(Chipperfield et al., 1998). Erle et al. (2000) recently

reported measurements of upper limits for OBrO that are

appreciably smaller than values observed by Renard et al.

(1997, 1998), indicating that one of the sets of observations is in error or that abundances of OBrO are highly

variable. The explanations for non-zero BrO mixing ratios

in darkness and possible detection of OBrO remain a

mystery.

Bromine Trends and Budget

During the Arctic winter of 1998/1999, vertical profiles of all known major organic bromine species were

measured between 9 and 28 km (Pfeilsticker et al., 2000)

(Figure 3-26). The expected profile for inorganic bromine

that was inferred from the source gases agrees well (i.e.,

differences are within the measurement uncertainties) with

a second profile that was estimated from spectroscopic

observations of BrO and a photochemical model estimate

of the BrO/Bry ratio (see Figure 3-26) (Pfeilsticker et al.,

2000). This result indicates that the budget of bromine and

its photochemistry in the lower stratosphere are reasonably well understood. For early 1999, the mixing ratio of

total bromine estimated at 25 km in air of 5.6-year mean

age was 18.4 (+1.8, -1.5) pptv based on organic precursor

measurements, and 21.5 ± 3.0 pptv from BrO measurements. This slight offset allows for the possibility of a

bromine influx of 3.1 (-2.9, +3.5) pptv from the troposphere to the stratosphere (Pfeilsticker et al., 2000).

Attempts to quantify temporal trends in inorganic

bromine in the stratosphere traditionally have been ham-
POLAR OZONE

3.36

0.5

1.5

2.5

3.5

1998 1999 2000

Date

Ny Ålesund, 79° N

Meas. AM

Meas. PM

Model AM

Model PM

0.5

1.5

2.5

3.5

1998 1999 2000

Date

Kiruna, 68° N

Meas. AM

Meas. PM

Model AM

Model PM

0.5

1.5

2.5

3.5

1998 1999 2000

Date BrO DSCD (90°-80°) [10 molec/cm ]

Arrival Heights, 78° S

Meas. AM

Meas. PM

Model AM

Model PM BrO DSCD (90°-80°) [10 molec/cm ] BrO DSCD (90°-80°) [10 molec/cm ]

Figure 3-25. Comparison of measured and modeled differential slant column densities (DSCD) of BrO over

Ny Ålesund, Spitzbergen (78.9°N), Kiruna, Sweden (67.9°N), and Arrival Heights, Antarctica (78°S). The differential slant columns were calculated from observations and model results obtained between solar zenith

angles of 90° and 80° for the morning (AM) and evening (PM). However, when a solar zenith angle of 80° was

not reached at high latitude (shaded regions), local noon values were used as a reference instead. Calculated

values of BrO are from the SLIMCAT model. Adapted from Sinnhuber et al. (2002).
POLAR OZONE

3.37

pered by the lack of long-term observations and the relatively small quantities (~10 pptv or less) of the bromine

species. However, a recent analysis of stratospheric measurements of BrO may shed some light on this issue, which

is important in the context of polar ozone loss because the

source of bromine to the stratosphere is expected to have

increased by nearly 30% over the past decade (Wamsley

et al., 1998).

Pfeilsticker et al. (2000) have taken the approach

of estimating the abundance of total inorganic bromine

(Bry) based on measurements of BrO and model calculations of the BrO/Bry ratio (Figure 1-8 in Chapter 1 of this

Assessment; see also Figure 2.12 of EC (2001)). They

rely on separate measurements of carbon dioxide (CO2)

or sulfur hexafluoride (SF6) and model simulations to

determine the age of air versus altitude. Their analysis of

stratospheric data obtained between 1996 and 2000,

reported in EC (2001), is consistent with a rate of increase

of total bromine of about 0.7 pptv/year, broadly consistent with the increase in tropospheric organic bromine

over this time period (Figure 1-8).

This result suggests that the contribution of

bromine to ozone loss in the polar regions has increased

faster than that of chlorine because of abundances of

bromine that continue to increase at a time when those of

chlorine are leveling off (see Chapter 1). Model studies

indicate that catalytic cycles involving BrO account for

as much as 60% (depending on abundances of ClO and

temperatures) of the total chemical loss of ozone in the

Arctic for cold winters (Chipperfield and Pyle, 1998). The

contribution of BrO reactions to the total loss of Antarctic

ozone is somewhat less than for the Arctic because of

lower temperatures and widespread denitrification in the

SH vortex. Considering the observed leveling off of chlorine sources gases, the relative role of bromine in polar

ozone loss will continue to increase until the upward trend

of bromine source gases reverses. However, since ozone

loss by the BrO cycle also depends on ClO, future major

declines in Cly are expected to lead to reductions in

chemical loss of polar ozone essentially independent of

changes to Bry (Chipperfield and Pyle, 1998).

3.2.4 The Polar Summer Lower

Stratosphere

In both hemispheres, the annual cycle of total ozone

has a strong decrease from the spring maximum to a minimum by mid-fall (e.g., Dobson, 1966; Dütsch, 1974;

Bowman and Krueger, 1985; see also Figure 3-1). Column abundance of ozone declines by ~35% at high

northern latitudes during summer (Toon et al., 1999; Lloyd

et al., 1999). The rate of chemical ozone destruction in

late spring/early summer is as large as in the winter polar

stratosphere. This large decline, and the fact that the

summer circulation is weak and quite zonally symmetric,

makes the summer period a good test of understanding.

Models typically have been unable to capture the full magnitude of the decline. To address this issue, the POLARIS

experiment was flown from Fairbanks, Alaska, with three

deployments during the early spring, summer, and early

fall of 1997. This mission examined the seasonal ozone

decrease using a complete payload of instruments aboard

the NASA ER-2 high-altitude aircraft (Newman et al.,

1999).

The summer lower stratosphere is mixed by waves

that penetrate in the presence of the weak summer westerlies (Wagner and Bowman, 2000). Orsolini (2001),

however, has shown that remnants of polar vortex air can

retain their identity until well into the summer in the lower

stratosphere. Rosenlof (1999) has studied the annual cycle

of ozone transport in high northern latitudes. She found

8

1 2

1 6

2 0

2 4

2 8

0 3 6 9 12 15 18 21 24

Organic Bry

Inorganic Bry

Kiruna, Sweden

February, 1999

MIXING RATIO (ppt) ALTITUDE (km) Figure 3-26. Vertical profile of total inorganic

bromine (labeled Inorganic Bry) based on the

observed falloff of brominated source gases (labeled

Organic Bry) and based on measured BrO and a

model calculation of the BrO/Bry ratio. Measurements were obtained in the Arctic vortex, above

Kiruna, Sweden (67.9°N), during the winter of

1998/1999. Adapted from Pfeilsticker et al. (2000).
POLAR OZONE

3.38

that the seasonal cycle in transport was an important contributor to the seasonal march of ozone at high latitudes

during middle to late summer, with the eddy contribution

to ozone reduction more than offsetting the ozone increase

by advection at this time. In early summer, transport is

weak and in situ photochemical destruction dominates the

ozone tendency. In contrast, Pierce et al. (1999) using a

Lagrangian model with HALOE data found that the transport term was important (see Appendix 3A for HALOE

data descriptions).

In situ measurements of nitrogen oxides (NOx), odd

hydrogen (HOx), and ClOx radical species (Fahey et al.,

2000) and long-lived tracers of stratospheric transport

(Toon et al., 1999) confirm the summertime loss of ozone

is due primarily to the gas-phase catalytic cycle:

(3a) NO + O3 Æ NO2 + O2

(3b) NO2 + O Æ NO + O2

Net: O + O3 Æ 2 O2 (3-3)

Fahey et al. (2000) calculated the ozone chemical

tendency based on measurements of radicals from the

major families and found that chemical processes dominate overall tendency in the midsummer. Although this

ozone loss process is generally understood (Brühl et al.,

1998), ozone abundances calculated using two- or threedimensional models tend to exceed observations for

high-latitude summer (e.g., Chipperfield, 1999). This discrepancy has been attributed to inadequacies in model

transport (e.g., Fahey and Ravishankara, 1999).

New measurements of NOx, HOx, and ClOx species

in the summer polar stratosphere have provided quantitative tests of our understanding of processes that regulate

the abundance of radicals in each family. Measurements

in the summer polar stratosphere are particularly useful

because heterogeneous reactions, normally the dominant

loss process for NOx, proceed at slower rates than gasphase loss reactions because of uninterrupted periods of

solar illumination that restrict the buildup of dinitrogen

pentoxide (N2O5) (e.g., Gao et al., 1999; Osterman et al.,

1999). In the sections that follow, we briefly assess recent

advances in our understanding of stratospheric photochemistry based on these observations.

3.2.4.1 SUMMERTIME NOX CHEMISTRY

Observations of NO, NO2, HNO3, and NOy

obtained in the summer polar stratosphere revealed higher

levels of NOx (relative to total NOy) and NO2 (relative to

HNO3) than could be accounted for by constrained photochemical box models using the JPL 97-4 (DeMore et al.,

1997) set of recommended kinetic parameters (Gao et al.,

1999; Jucks et al., 1999; Osterman et al., 1999; Cohen et

al., 2000; Perkins et al., 2001) (Figure 3-27). During polar

summer, production and loss of NOx are regulated primarily by the OH + NO2 and OH + HNO3 reactions for a

broad range of altitudes (e.g., Osterman et al., 1999). New

laboratory data for OH + NO2 (Dransfield et al., 1999;

Brown et al., 1999a) and OH + HNO3 (Brown et al.,

1999b) led to a re-evaluation of the rate constant for both

of these reactions in the JPL 00-3 compendium (Sander et

al., 2000). Use of the JPL 00-3 kinetic parameters significantly improves the agreement between measured and

modeled ratios of NOx/NOy and NO2/HNO3 (Gao et al.,

1999; Jucks et al., 1999; Osterman et al., 1999; Cohen et

al., 2000; Perkins et al., 2001; Salawitch et al., 2002b)

(Figure 3-27).

It has also been proposed that the reaction of OH +

NO2 produces HOONO (pernitrous acid) in addition to

HNO3 (nitric acid) (McGrath and Rowland, 1994; Matheu

and Green, 2000; Golden and Smith, 2000, and references

therein). Recent spectroscopic observations indicate a

yield for HOONO of about 5% at 253 K and 26 hPa

(Nizkorodov and Wennberg, 2002). The formation of the

HOONO isomer is also supported by several recent kinetic

studies (Donahue et al., 2001; Hippler et al., 2002).

Implications of HOONO production have yet to be fully

explored, but box model calculations of Golden and Smith

(2000) indicate this process leads to small increases in

calculated NOx, and small decreases in calculated ozone,

for the high-latitude summer stratosphere.

Simultaneous measurements of ozone, NO, NO2,

ClO, and HO2 during polar summer provide a stringent

test of our understanding of the rapid photochemistry

linking NO to NO2 (Del Negro et al., 1999). Photolysis

rates for NO2 (JNO2) inferred from the chemical measurements are in excellent agreement with values calculated

with radiative models and those measured with a spectroradiometer (Del Negro et al., 1999). Recently, this comparison has been extended to larger SZAs of 80° to 93°

(Gao et al., 2001a). Values of JNO2 derived from the in

situ chemical measurements agree well (differences

<11%) with results from a multiple-scattering actinic flux

model. The linearity of the correlation between these two

computations of JNO2 over the SZA range 80° to 93°

demonstrates that the model scattering calculation is

accurate for twilight conditions (Gao et al., 2001a).

Loss of ozone by NOx chemistry in the summer

polar regions may become more important in the future

because of rising levels of NO2 (Liley et al., 2000;

McLinden et al., 2001). Nitrous oxide (N2O), the source

gas for NOx and NOy, is rising at about 3% per decade

(e.g., McLinden et al., 2001). Interestingly, concentra-
POLAR OZONE

3.39

tions of NO2 at SH midlatitudes have been observed to be

increasing at a faster rate of 5% per decade (Liley et al.,

2000). This increase has been interpreted as being due to

rising N2O as well as declining levels of ozone, which

alter the NO/NO2 partitioning and the diurnal variation of

NOx (McLinden et al., 2001); see also Chapter 4.

3.2.4.2 SUMMERTIME HOX CHEMISTRY

Measurements of OH and hydroperoxyl radical

(HO2) in the high-latitude stratosphere during late spring

provide an important test of our understanding of HOx

sources throughout the day because abundances of these

species are nearly in photochemical steady state under the

slowly varying SZA conditions up to 93° (Wennberg et

al., 1999). Observed abundances of OH and HO2 in the

lower stratosphere significantly exceed those from standard model calculations for SZA > 80° (Wennberg et al.,

1999; Salawitch et al., 2002b) (Figure 3-28).

An important change in the JPL 00-3 kinetics

evaluation, relative to the JPL 97-4 evaluation, is that

the recommended reaction probability for BrNO3 hydrolysis dropped from 0.8 to about 0.2 for conditions of the

high-latitude, springtime lower stratosphere. Consequently,

a model using JPL 00-3 kinetics (Figure 3-28) does not

predict the observed “morning rise” of HO2; the calculated nighttime buildup of HOBr, which releases HOx at

sunrise, is limited because of this rate change (Salawitch

et al., 2002b). The OH and HO2 measurements suggest

the presence of a photolytic source of HOx that operates

more efficiently than known HOx sources during twilight and that operates at an essentially constant rate for

80° < SZA < 93° (e.g., this photolytic process must occur

Figure 3-27. (a) MkIV observations (red points with error bars) of the NO2/HNO3 ratio at sunrise obtained

during a balloon flight in the Arctic stratosphere on 8 May 1997. Constrained photochemical steady-state

model calculations of the ratio, for the solar zenith angle of the observations, are shown for kinetic parameters

from the JPL 00-3 (Sander et al., 2000) evaluation (solid line) and from the JPL 97-4 (DeMore et al., 1997)

evaluation (dashed line). Error bars on the observations represent 1s precision, based on considerations

such as residuals in the spectral fitting, combined in quadrature with spectroscopic uncertainties. (b)

Percentage difference between measured NO2/HNO3 and values from the two model calculations. The red

dotted lines denote the 1s measurement uncertainty. After Osterman et al. (1999) and Salawitch et al. (2002b).
POLAR OZONE

3.40

longward of 650 nm, which remains optically thin even

for SZAs near 93°) (Wennberg et al., 1999).

Including a photolytic pathway for photolysis of

HO2NO2 (peroxynitric acid) via excitation of purely vibration modes longward of 760 nm (the near IR) based on

recent laboratory measurements of cross sections and

quantum yields (Roehl et al., 2001), a process first suggested by Donaldson et al. (1997), leads to significant

improvements in measured and modeled HOx near twilight due to the rapid photolysis of HO2NO2 (Wennberg et

al., 1999; Salawitch et al., 2002b) (Figure 3-28). The nearIR photolysis of HO2NO2 reduces calculated levels of this

species, resolving a large discrepancy between standard

model calculations and observations of HO2NO2 at high

latitudes (Salawitch et al., 2002b). The lower calculated

abundance of HO2NO2 also reduces the efficiency of the

OH + HO2NO2 sink of HOx, resulting in rather large (e.g.,

20 to 60%) increases in 24 hour averaged HOx for the

high-latitude spring lower stratosphere and upper troposphere compared with models that neglect this process

(Salawitch et al., 2002b). Nonetheless, the observed

“morning rise” of HO2 is still underestimated by a model

including near-IR photolysis of HO2NO2. This discrepancy might be due to the actual rate of BrNO3 hydrolysis

proceeding at a faster rate than the JPL 00-3 recommendation, the presence of other twilight photolytic sources

of HO2 besides HOBr and HO2NO2, less than 100% efficiency for production of H2O from the OH + HO2NO2

reaction, or a systematic error in the measurements of OH

and HO2 (Wennberg et al., 1999; Salawitch et al., 2002b).

Balloonborne observations of OH, HO2, H2O, and

ozone obtained over Fairbanks, Alaska, also suggest

important gaps in our understanding of several HOx reactions (Jucks et al., 1998). Discrepancies between measured and modeled abundances of OH and HO2 are reduced

with a ~25% downward adjustment of the ratio of the rate

constants for atomic oxygen (O) + HO2 and O + OH and

either a 25% reduction to the rate constant for OH + HO2

(the primary HOx sink) or a 25% increase in the HOx production rate (Jucks et al., 1998). These modifications are

within the uncertainties of the laboratory measurements

and are consistent with the results of HOx model/measurement studies discussed in the previous Assessment.

The new insight provided by the simultaneous observations of OH and HO2 is that the required modifications to

the rate constants appear not to appreciably affect odd

oxygen production rates in the upper stratosphere (Jucks

et al., 1998).

3.2.4.3 SUMMERTIME ClY CHEMISTRY

The first in situ observations of ClNO3 were

obtained in the high-latitude stratosphere during the

summer of 1997 (Stimpfle et al., 1999). These observations are in good agreement with values of ClNO3 determined using a photochemical steady-state relation constrained by simultaneous observations of [ClO] and [NO2]

(the ratio of measured to modeled ClNO3 is 1.15 ± 0.36).

These results, together with a study that used balloonborne remote measurements (Sen et al., 1999), confirm

the photochemical mechanism by which abundances of

NOx regulate the abundance of ClO in regions of the stratosphere that are NOx-limited (i.e., mixing ratio of NOx >

mixing ratio of ClOx).

Simultaneous observations of ClNO3 and HCl from

a balloon (Sen et al., 1999) and the ER-2 aircraft (Voss et

al., 2001) test our understanding of the kinetic processes

that regulate the partitioning within the inorganic chlorine

(Cly) family. Model calculations using JPL 00-3 recommendations agree extremely well with the balloonborne

remote observations of both ClNO3 and HCl (Figure 3-

29), whereas the aircraft in situ measurements of the ratio

[ClNO3]/[HCl] are ~55 to 60% lower than values based

on a steady-state calculation (Voss et al., 2001). This disNear IR

JPL00-3

Figure 3-28. Observations of HO2 obtained on the

morning of 30 April 1997 and the afternoon of 9 May

1997 near 64°N in the lower stratosphere (Wennberg

et al., 1999) compared with constrained photochemical model simulations using three sets of kinetic

parameters: (1) JPL 00-3 (Sander et al., 2000) (red

dashed lines); (2) allowing for near-IR photolysis of

HO2NO2 based on cross sections and quantum

yields from Roehl et al. (2001) (green dashed lines);

and (3) allowing for near-IR photolysis of HO2NO2

and for a reaction probability of 0.8 for BrNO3 hydrolysis (blue solid lines). After Salawitch et al. (2002b).
crepancy has not been resolved and is the subject of

ongoing investigations.

Simultaneous observations of the major inorganic

and organic chlorine species in the summer polar stratosphere provide a test of the chlorine budget. The observed

increase in the inorganic chlorine content (e.g., HCl +

ClNO3 + HOCl + ClO) of stratospheric air with decreasing

N2O, from both the ER-2 and MkIV, agrees well (differences less than ~10%) with estimates of inorganic chlorine based on the observed disappearance of organic

source molecules (Sen et al., 1999; Bonne et al., 2000)

(Figure 3-29). However, the ER-2-based estimate of inorganic chlorine is ~15% less than the MkIV estimate for

mixing ratios of N2O below about 175 ppbv. This offset

is probably due to differences in the measurements of

ClNO3 noted in the previous paragraph. These results

extend the conclusions of the previous Assessment

regarding the good quantitative link between abundances

of inorganic chlorine species and their halogen sources to

a new region of the atmosphere (i.e., polar summer) and

to a new class of observations (i.e., in situ). These findings are particularly relevant for the discussion of the chlorine budget for the winter polar stratosphere (Section

3.2.3.1).

3.3 QUANTIFICATION OF POLAR OZONE

LOSS: OBSERVATIONS AND MODELS

Quantification of the degree of chemical ozone loss

in the polar stratosphere is hampered by the pronounced

dynamically induced variability of the ozone layer in these

regions in winter/spring. Precise quantification of the

chemically induced contribution to observed changes in

the ozone abundance are particularly difficult in the Arctic

stratosphere, where the degree of ozone loss is smaller and

the dynamic activity is more pronounced than in the

Antarctic. Over the last decade a number of approaches

have been developed to overcome these difficulties. Major

challenges have been to (1) assess how reliable the results

of these approaches are and, hence, how precise current

estimates of the degree of Arctic ozone losses are, and (2)

determine whether the degrees of ozone losses calculated

by up-to-date chemical models agree with the observational results within the combined uncertainties of the

models and the observations. The latter question is crucial

to assess our current ability to make projections of future

polar ozone losses in a potentially colder stratosphere.

Severe Arctic ozone loss has been reported for

some recent cold Arctic winters. No significant loss was

found during warmer winters. The large Arctic ozone

losses result from increased levels of radical halogen

species that result from heterogeneous chemical processing on the surface of polar stratospheric clouds

(WMO, 1999; cf. also Sections 3.2 and 3.3.2 of this

Assessment). Currently about 80 to 85% of the stratospheric chlorine is provided by the decomposition of

anthropogenic organic chlorinated species (Zander et al.,

POLAR OZONE

3.41

Figure 3-29. Measurements of chlorine species

from balloon and in situ instruments during

POLARIS. MkIV balloonborne measurements of the

volume mixing ratio of HCl (green open triangles)

and ClNO3 (red solid squares) versus N2O observed

over Fairbanks, Alaska (64.8°N), on 8 May 1997 are

shown, along with calculated values of HCl and

ClNO3 (solid black lines) found using a constrained

photochemical model (Sen et al., 1999, updated for

JPL 00-3 (Sander et al., 2000) kinetics). The dashed

blue line shows an estimate of Cly based on ER-2

measurements of the organic source molecules

during POLARIS (Sen et al., 1999). The sum HCl +

ClNO3 + HOCl measured by MkIV plus calculated

ClO (aqua solid dots) is also shown; calculated ClO

makes a negligible contribution to this sum except

for the highest altitudes (e.g., N2O < 50 ppb)

because the MkIV observations are obtained at twilight. Error bars for the MkIV measurements represent 1s precision. ER-2 measurements of the sum

HCl + ClNO3 + ClO are shown by the black dots

(Bonne et al., 2000). Because the ER-2 observations are obtained at altitudes below ~20 km,

species such as HOCl, Cl, etc., are estimated to

make a negligible (e.g., <20 ppt) contribution to the

sum and are therefore neglected. Finally, the arrow

in the right margin denotes the total chlorine content,

3.59 ppb, of the entire suite of tropospheric organic

compounds for 1996 (WMO, 1999, Table 1-2).
1992, 1996; Gunson et al., 1994; Russell et al., 1996; Sen

et al., 1999). Because the relevant ozone destruction

cycles are linear to quadratic in the concentration of active

chlorine, by far the largest fraction of the ozone losses

discussed in this section is of anthropogenic origin.

However, no quantitative study exists to precisely quantify the small degree of ozone loss that would be expected

in a cold Arctic winter for natural levels of halogens in

the stratosphere.

3.3.1 Approaches to Quantify Chemically

Induced Ozone Loss in the Arctic

Two principal techniques are currently used to

quantify chemically induced ozone losses in the Arctic:

(1) Studies that take into account the effect of transport

explicitly by using transport calculations based on

meteorological analyses.

(2) Studies that allow for transport effects implicitly by

using the relation between ozone and a long-lived

chemical tracer.

In the following, approaches that have been used

in a consistent way for several winters are briefly assessed.

3.3.1.1 APPROACHES THAT USE EXPLICIT TRANSPORT

CALCULATIONS

Bulk Advection

In this section, approaches are assessed that use

explicit transport calculations to advect bulk quantities

like vortex averages or gridded ozone fields and compare

these with later measurements of ozone, a concept first

published by Manney et al. (1994).

The “vortex average” technique involves analysis

of the temporal evolution of the mean profile of ozone

within the polar vortex on surfaces of potential temperature. Ozone measurements by ozonesondes (e.g.,

Knudsen et al., 1998; Rex et al., 1998; Lucic et al., 1999;

Tzvetkova et al., 2002) or a remote sensing instrument,

e.g., the POAM II satellite instrument (Bevilacqua et al.,

1997) have been analyzed with this approach. The

boundary of the polar vortex is usually defined by isolines

of potential vorticity. Potential temperature and potential

vorticity are conserved quantities in the polar stratosphere

over time scales of a couple of weeks, so this coordinate

system largely eliminates variability due to rapid and

reversible dynamics. The slow irreversible descent of air

across surfaces of potential temperature is usually

accounted for by a diabatic correction calculated from

descent rates using a radiative transfer model that is based

on temperatures from a meteorological assimilation

system. In the absence of mixing across the vortex edge,

changes in ozone can be attributed to chemical loss. The

largest uncertainties connected with this approach are (1)

possible mixing across the vortex edge, which may impact

the average ozone abundance inside the polar vortex, (2)

any time-varying bias in the sampling of the vortex, which

may lead to changes in the derived vortex average, since

ozone is not uniform within the vortex, and (3) uncertainties in the calculated diabatic corrections. Knudsen et al.

(1998) used an approach based on domain-filling trajectory calculations to estimate the effect of mixing across

the vortex edge for the Arctic winter of 1996/1997. For

that winter, at the altitude of the maximum loss, they found

an insignificant impact of mixing on the ozone loss

derived from the vortex average technique.

In the “transport model” approach, a 3-D CTM is

initialized with ozone observations during the early winter

period. The model advects ozone passively, i.e., without

chemical conversion, throughout the winter, using analyzed winds and, depending on the vertical transport

scheme in the model, also temperatures. Ozone measurements throughout the winter are compared with the

passively advected ozone. Evolving deficits between

observed ozone and the model passive ozone indicate

chemical loss. This approach has been used with the

Reactive Processes Ruling the Ozone Budget in the

Stratosphere (REPROBUS) model using data from the

POAM satellite instruments and from the ground-based

network of Système d’Analyse par Observation Zénithale

(SAOZ) instruments (e.g., Lefèvre et al., 1998; Goutail et

al., 1999; Deniel et al., 2000). Hansen et al. (1997),

Guirlet et al. (2000), and Sinnhuber et al. (2000) have used

this approach with the SLIMCAT model, using ozone lidar

data, SAOZ data, and ozonesonde data, respectively. This

approach relies on the assumption that the model transport scheme realistically represents the transport of air

over a time period of several months. But when the

derived ozone loss is averaged over the polar vortex, as is

the case in many studies, the approach only depends on a

correct representation of average transport properties like,

e.g., vortex-averaged vertical subsidence and average

rates of exchange of air across the vortex edge. Possible

systematic errors in these average properties are the largest

source of uncertainty in the transport model approach.

Another source of concern is the initialization of the model

ozone with data from a different type of instrument (usually measurements of the HALOE or MLS satellite instruments; see Appendix 3A for satellite data descriptions)

than is used for the ozone observations later on. Systematic discrepancies between the different instruments

make altitude-dependent correction factors necessary;

these are often on the order of 5%, which is a substantial

POLAR OZONE

3.42
fraction of the ozone changes observed over the course of

warmer Arctic winters. The method is sensitive to the vertical variation of the correction, because early-winter

measurements from one instrument at higher altitude are

compared with late-winter measurements from another

instrument at lower altitude.

Manney et al. (e.g., 1995a, b, 1996a, b, 1997) used

a similar approach based on “trajectory ensemble” calculations to analyze ozone data of the MLS on the UARS

satellite (see Appendix 3A for MLS data description).

Trajectory calculations are started at all points on the

gridded MLS data and are run forward in time for a few

weeks. Succeeding MLS measurements are interpolated

to the locations of the trajectories. The differences in

ozone are attributed to chemistry. The results are stated

as vortex-averaged ozone loss or are analyzed versus

equivalent latitude, PV, etc. During the UARS northlooking yaw cycle, continuous time series of ozone loss

can be derived, whereas only accumulated losses are

stated across the south-looking yaw cycles, when measurements in the Arctic are not available. Schoeberl et al.

(2002) advected ozone, using trajectory ensemble calculations initialized from early-winter ozone observations,

over the course of the winter and then compared this early

winter ozone with late-winter ozone observations. Again

the advected and the observed ozone fields are averaged

over the polar vortex before they are compared, thus

reducing the sensitivity of the approach to transport features of individual air masses. The trajectory ensemble

approach uses 3-D trajectory calculations that are several

months long for the advection. Using trajectories instead

of a grid-point advection scheme eliminates any potential

bias due to numerical diffusion of the advection scheme.

The largest uncertainty of the trajectory ensemble method

comes from possible systematic biases in the long-term

trajectory calculations, e.g., due to possible uncertainties

in the vertical transport, which is based on calculated diabatic descent rates.

Lagrangian Ozone Measurements (Match)

“Match” is a Lagrangian technique to determine

the rate of chemical ozone loss. In active Match campaigns, ozonesonde launches from a large network of

about 35 ground stations are coordinated in real time to

probe individual air masses twice over an interval of a few

days (so-called “match events,” e.g., von der Gathen et

al., 1995; Rex et al., 1997, 1998, 1999b, 2002a). The coordination is based on calculations of air parcel trajectories

that allow for diabatic descent. Several hundreds to more

than a thousand ozonesondes are launched in a Match campaign. The initial launch subsequently results in several

match events with later sondes at different altitudes. The

coordination results in hundreds to thousands of match

events per winter. Chemical ozone loss rates are derived

from a statistical analysis of subsets of match events from

a certain time period (typically 14 days long) and altitude

region (typically 20 K broad) by calculating linear regressions of the difference in ozone between both measurements and the sunlit time that the air mass encountered.

The overall ozone loss during the winter is calculated by

accumulating the measured loss rates. Sasano et al. (2000)

and Terao et al. (2002) have used a similar approach to

analyze the ozone measurements from the ILAS satellite

instrument (see Appendix 3A for ILAS data descriptions).

The largest uncertainties in this method are possible systematic errors in the trajectory calculations,

including the calculated diabatic subsidence rates.

Individual Match events depend on a correct representation of the motion of individual air masses by the trajectories. Therefore, the length of the trajectory calculations

used in Match is limited to 10 days, and the majority of

the Match events rely on ~5-7 days trajectories. But ozone

loss rates are calculated in a statistical process from subsets of several tens of Match events, so that the derived

average ozone loss rates are only sensitive to systematic

trajectory errors, e.g., possible systematic biases in the

calculated diabatic descent rates. A statistical analysis of

the data shows that observed ozone losses occur exclusively during sunlit periods along the trajectories. This

suggests that any systematic biases in the trajectory calculations, if present, are so small that they do not significantly affect the derived ozone loss rates (e.g., Rex et al.,

1998, 1999b, 2002b).

3.3.1.2 APPROACHES THAT USE THE RELATION OF

OZONE TO AN INERT TRACER

In midwinter, ozone abundances inside the polar

vortex show a relatively compact relation to abundances

of many long-lived tracer species such as nitrous oxide

(N2O) or methane (CH4). In the absence of mixing, any

reduction of ozone versus an inert tracer indicates chemical loss of ozone. This approach was first used by Proffitt

et al. (1990). More recently, early-winter and late-winter

measurements of ozone, N2O, CH4, and HF by the

HALOE satellite instrument have been used to identify

chemical loss of ozone during various winters (e.g.,

Müller et al., 1996, 1997; see also Appendix 3A for satellite data descriptions). The evolution of the ozone-versusN2O relation through the winter of 1999/2000 was studied

using data from in situ instruments onboard two balloonborne platforms and the NASA ER-2 high-altitude aircraft (Salawitch et al., 2002a; Richard et al., 2001).

POLAR OZONE

3.43
Two fundamental issues affect the validity of the

tracer relation approach. First, the existence of a compact

universal relation between ozone and inert tracers has been

questioned, and, second, the impact of mixing on results

from studies of the ozone-versus-tracer relation has been

discussed. Here we assess the current status of the discussion on these two points.

During polar summer the chemical lifetime of

ozone in the middle and upper stratosphere is comparable

with or shorter than transport time scales, so the ozonetracer relationship in the Arctic stratosphere in fall is not

expected to be compact and universal (Plumb and Ko,

1992). But in fall, the lifetime of ozone in the middle and

lower stratosphere gets sufficiently long, and it can be

expected that mixing within the early polar vortex leads

to compacting of the ozone-tracer relations inside the

vortex. These slowly evolving, more-compact relations

inside the vortex are quite different from the extra-vortex

relations, with less ozone inside at a given tracer level

than outside. The degree of compactness and the definition of this inner vortex “early-winter relation” is critical

for the validity of the tracer relation approach. Richard et

al. (2001) and Salawitch et al. (2002a) used initial ozoneversus-N2O reference relations measured well inside the

polar vortex in midwinter (December/early-January).

They showed that these relations were sufficiently compact and representative of initial conditions inside the

polar vortex, an essential condition for the validity of the

approach. But observations of the ozone-versus-methane

relation measured well inside the Arctic vortex in midwinter for 1999/2000 were significantly lower (ozone differences of ~1.5 ppmv for values of CH4 = 0.5 ppmv) than

the suite of initial HALOE-based reference relations

(Müller et al., 1999, and references therein) measured in

the vortex edge region in October for earlier Arctic winters (Salawitch et al., 2002a). The validity of these earlywinter HALOE reference relations is currently the subject of numerous ongoing investigations (see Appendix

3A for satellite data descriptions). Generally results from

the tracer relation approach are more reliable when the

initial reference relation is measured late (i.e., December/

early-January) and deep inside the vortex.

The second issue brought up regarding the validity

of tracer relation studies is related to mixing. These considerations apply also for studies that assess the degree of

denitrification based on changes in the otherwise very

compact relation between NOy and N2O. If the relation

between two tracers is curved, as is the case for these

species, the results of the tracer relation approach can be

compromised by mixing between air masses that are

widely separated in tracer space (e.g., Waugh et al., 1997;

Michelsen et al., 1998; Rex et al., 1999a; Plumb et al.,

2000; Ray et al., 2002). Isentropic mixing across the edge

of the polar vortex or mixing of air masses inside the

vortex that underwent different descent during the winter

are examples for such long-range mixing in tracer space.

Furthermore, the ozone/tracer relations inside the vortex

are different from the relations outside the polar vortex.

After substantial ozone loss, these differences can be very

pronounced (several ppmv). Hence, any mixing across

the vortex edge directly impacts the ozone/tracer relation

inside the vortex and could represent a potential source of

uncertainty for the tracer relation approach.

One can attempt to distinguish mixing and chemical ozone loss (or denitrification) by using simultaneous

measurements of two long-lived tracers (e.g., CH4 and

N2O) to estimate the impact of mixing (Rex et al., 1999a).

However, this method is dependent on the assumption

that there had been a single mixing event after the bulk

of the descent and is not applicable if there is intermittent mixing at descent (Plumb et al., 2000). A more general approach is the use of a linear combination of several long-lived tracers to form an artificial tracer that has

a compact and linear relationship with ozone (or NOy)

(Esler and Waugh, 2002). Because of its linearity, this

artificial relation is unaffected by mixing within the

vortex, so deviations from this relationship can be more

directly attributed to chemical ozone loss (denitrification). But mixing across the edge of the vortex edge is

still a source of uncertainty, because the outside vortex

relations are often different from the inside relations (for

ozone versus N2O this is always the case; for NOy versus

N2O this is the case after denitrification inside of the

vortex). These attempts to correct for the impact of

mixing in tracer relation studies were mainly focused on

studies that used the relation of NOy versus N2O to assess

denitrification.

Salawitch et al. (2002a) noted that considerations

that are valid for the NOy-versus-N2O relation (e.g., the

relation c2 versus c1 in Plumb et al. (2000) resembles

NOy versus N2O) should not be applied to the interpretation of the ozone-versus-N2O relation in the vortex,

because ozone mixing ratios, unlike NOy, do not approach

zero at the top of the vortex because of the influence of

photochemistry at 40 km. Hence, the curvature of the

ozone-versus-N2O relation for low N2O (i.e., N2O between

10 and 40 ppbv) is much less pronounced than that of NOy

versus N2O, and mixing cannot lead to the observed

changes in the ozone tracer relations that have been

observed e.g., by HALOE. More quantitatively, simultaneous measurements of multiple long-lived tracers have

been used to argue that the impact of mixing on estimates

of chemical ozone loss by the tracer relation approach for

the Arctic winter of 1999/2000 was negligible. Based on

POLAR OZONE

3.44
the temporal evolution of CO2, CFC-11 (CCl3F), N2O, and

ozone within the vortex, studies of Richard et al. (2001)

and Salawitch et al. (2002a) demonstrated that the vast

majority of the observed changes in the O3/N2O relations

were due to chemistry and could not have been caused by

dynamics. The isolation of the Arctic vortex for that

winter was also noted by a multivariate analysis of the

time evolution of nearly a dozen tracers with varying lifetimes (Ray et al., 2002). Rex et al. (2002a) showed that

during January to March 2000, any mixing across the

vortex edge would have led to an underestimation of the

ozone loss by tracer relation studies, and therefore the

results of tracer relation studies that rely on initial relations from early January can be regarded as conservative

estimates of the loss. These results from the Arctic winter

of 1999/2000 support the validity of the tracer relation

approach, provided the reference relation is defined in

midwinter, but it is currently not clear whether these

results can be applied to other, more dynamically active

winters (e.g., Salawitch et al., 2002a).

3.3.2 Arctic Ozone Loss During the Last

Decade

Since the previous Assessment (WMO, 1999), a

number of approaches to quantify the degree of chemical

ozone loss in the Arctic have been used in a consistent

way for several winters during the 1990s. Because of

these long-term efforts, the extent and the variability of

Arctic ozone losses are now well characterized for the last

10 years. Several techniques have revealed a large interannual variability of chemical ozone losses in the Arctic.

3.3.2.1 OZONE LOSS RATES NEAR THE MAXIMUM OF

THE OZONE CONCENTRATION

The Match approach was used consistently over

the last decade to study the evolution of chemical ozone

loss at about 475 K potential temperature (~19 km altitude). At polar latitudes this level is close to the maximum concentration of ozone in the ozone layer. Figure

3-30 shows the measured ozone loss rates for the winters

1991/1992 to 2000/2001 (compilation of Match results

based on Rex et al. (1997, 1998, 1999b, 2002a) and Schulz

et al. (2000, 2001)). Little or no significant ozone loss

was observed in 1997/1998 and 1998/1999. The accumulated ozone loss during the winter was particularly large

in 1995/1996 and 1999/2000, when relatively large loss

rates were sustained for extended periods of time. The

chemical ozone loss rate in the Arctic stratosphere is

clearly controlled by temperature. Blue shaded areas in

Figure 3-30 indicate the geographical areas (APSC) where

temperatures have been below TPSC (the NAT equilibrium

temperature based on 5-ppmv water vapor and an average

HNO3 profile based on measurements; TPSC is a convenient threshold that roughly indicates the onset of rapid heterogeneous chemistry in the stratosphere, independent of

the actual composition of the PSCs). All periods of rapid

chemical ozone loss in Figure 3-30 are associated with

preceding large values of APSC. No significant chemical

loss of ozone was observed in warm winters, when TPSC

was not reached or only barely reached. 

3.3.2.2 VERTICAL PROFILES OF OZONE LOSS

The determination of vertical profiles of ozone loss

requires quantification of ozone losses over a broad altitude range. Ozone loss observations are typically most

reliable at levels around 475 K. Below 400 K or above

550 K the uncertainty of observational studies is typically

significantly larger. The problems at lower levels are that

(1) a strong vertical gradient in the average ozone mixing

ratio profile largely amplifies the uncertainty introduced

by diabatic descent, and (2) larger dynamical activity via

synoptic-scale waves makes explicit transport calculations less reliable and leads to larger degrees of mixing,

which is problematic for all approaches. Above 550 K,

average poleward motion and influx into the polar vortex,

followed by mixing, is the main problem. The strong diabatic descent at higher altitudes causes additional uncertainty only in winters when the vertical gradient in the

average ozone mixing ratio profile is significant at these

levels, which is not always the case (e.g., compare

1998/1999 with 1997/1998 in Figure 3-31).

Figure 3-31 shows ozone losses derived from the

vortex average approach in the vertical region between

Q = 360 and 570 K (update from Rex et al. (2002a) for

various winters). Results at the lowest and highest levels

shown are less reliable. It is based on several hundred

ozonesonde measurements per winter inside the vortex

from a network of about 35 sounding stations. The vortexaveraged ozone profiles have been plotted against the

“spring-equivalent potential temperature” (eQ), which is

the potential temperature that a given air mass reached at

the end of March due to diabatic subsidence. Through

use of eQ, which is a conserved quantity, diabatic effects

are accounted for. In the absence of mixing across the

vortex edge, any change in the vortex-averaged ozone

versus eQ profile indicates chemical loss of ozone. In

Figure 3-31, the large interannual variability of the ozone

loss stands out. During the winter of 1998/1999, no significant loss of ozone was found at any part of the profile.

In contrast, the loss of ozone in 1999/2000 exceeded 70%

in a ~1-km-thick region centered around 460 K. This local

POLAR OZONE

3.45
POLAR OZONE

3.46 Area of possible PSC existence [ 106 km2 ]

15

0

1995

-80

-40 Ozone depletion rate [ ppbv / day ] 0

1992 1993

1997

1999

1994

preliminary

2001

1996

2000

1998

0 30 90 60 0 30 90 60

Date [ day of year ]

-80

-40

0

-80

-40

0

-80

-40

0

-80

-40

0

15

0

15

0

15

0

15

0

Figure 3-30. Ozone loss rates from Match at 475 K for various winters (squares) and APSC at 475 K (blue

shaded areas). Red squares indicate 10 matches or more; white squares indiciate fewer than 10 matches.

Error bars are the 1s errors of the fit of ozone differences (matches) as a function of sunlit time in the time

period considered.
POLAR OZONE

3.47

550

500

450

400 adjusted potential temperature [ K ]

550

500

450

400

550

500

450

400

550

500

450

400

550

500

450

400

 2.4 10 km

(b) 1992/1993

 1.2 10 km

(a) 1991/1992

ozone [ ppmv ]

0 4 2 3 0 4 2 3

[APSC] = 5.4 10 km

(i) 1999/2000

 0.6 10 km

(g) 1997/1998

 3.3 10 km

(d) 1994/1995

 1.1 10 km

(c) 1993/1994

 5.0 10 km

(e) 1995/1996

 2.3 10 km

(f) 1996/1997

 0.0 10 km

(h) 1998/1999

 ] = 1.5 10 km

(j) 2000/2001

day of the year 0 30 60 90

[APSC

[APSC ] = [APSC ] =

[APSC ] = [APSC ] =

[APSC ] = [APSC ] =

[APSC ] = [APSC ] =

Figure 3-31. The evolution of the average ozone profile inside the Arctic polar vortex in spring-equivalent

potential temperature (eQ) coordinates for various years; see text for a definition of eQ. The data from all

ozonesonde measurements inside the vortex have been averaged into 10-day bins, centered around the day

indicated by the color scale. The profiles typically rely on 20 to 50 individual measurements that, in most

cases, sample the PV-space homogeneously. The average values for APSC ([APSC]) from 16 December to 31

March of each winter, between 400 and 550 K, are also given in each panel.
loss is slightly more than in any previous Arctic winter,

with 64% local loss in winter of 1995/1996 as the previous record. However, in 1995/1996, ozone loss occurred

over a broader vertical region (e.g., ozone loss of more

than 1 ppmv occurred between ~390 to 530 K in 1995/

1996, compared with ~420 to 510 K in 1999/2000), and

therefore the vertically integrated losses in both years are

comparable.

The average value of APSC (denoted as [APSC]), averaged from mid-December to end of March between 400

and 550 K potential temperature, is given in Figure 3-31.

Figure 3-32 shows the relation between [APSC] and the

average accumulated ozone loss between 400 and 550 K

eQ. A surprisingly close quantitative relation between

both quantities suggests that the chemical loss of ozone in

the Arctic stratosphere in a given winter correlates

strongly with the parameter [APSC]. The compactness of

the relation shown in Figure 3-32 is currently not fully

understood. Many mechanisms that influence Arctic

ozone loss, e.g., the degree of denitrification, should correlate well with the parameter [APSC]. But other meteorological parameters, like the timing of low-temperature

periods, which also impact the overall degree of ozone

loss, are not expected to correlate with [APSC]. Figure 3-

32 suggests that the effect of variability in these parameters has been relatively limited during the past decade. To

reproduce the empirical relation shown in Figure 3-32 is

a major challenge for global chemistry transport models.

The ability of models to reproduce the slope of the relation shown in Figure 3-32 is crucial, if models are to be

used to predict the impact of climate changes on future

ozone losses.

3.3.2.3 EFFECT OF OZONE LOSS ON THE TOTAL

OZONE COLUMN

Estimating the total column loss of ozone also

requires a good quantification of ozone loss in a broad

vertical region. The region above 550 K is of less concern for ozone column loss estimates because it contributes little to the total column amount of ozone, due to

the small ozone concentrations at these altitudes. But

uncertainties in the ozone loss estimates at altitudes below

400 K make ozone column loss estimates generally less

reliable than estimates of local ozone losses near 19-20

km. The ozone column loss has been estimated for all

winters since 1993/1994 with the transport model

approach. Figure 3-33 shows the difference between

ozone columns as measured by the SAOZ UV-visible network in the Arctic (Ny Ålesund, Thule, Scoresbysund,

Sodankylä, Salekhard, Zhigansk, and Harestua) and the

column of passive ozone in REPROBUS (again, initialized with POAM measurements in early winter) above

these stations. In many winters, large deficits of observed

ozone compared with passively advected ozone have been

observed. These deficits are attributed to chemical loss

POLAR OZONE

O loss

(a) 1991/1992 1994/1995 1995/1996 1996/1997 1997/1998 1998/1999 1999/2000 [A ] PSC (10 km [A ] PSC (106 km2)

0 2 3456 1992/1993 1993/1994246

(b) O3loss [ppmv]

1.5

0.5

[A ] PSC

Figure 3-32. (a) Histogram of average values of APSC ([APSC]; black bars) and the average value for chemical

loss of ozone from mid-January to the end of March, between 400 and 550 K (red bars). When the period of

rapid ozone loss overlaps with the vortex break up, the vortex average approach cannot be used to estimate

the overall loss of ozone. Those years (1994, 2001) have been omitted. (b) Scatter diagram of average ozone

loss versus [APSC].

3.48
POLAR OZONE

3.49

-40

-20

0

93/94

-40

-20

0

94/95

-40

-20

0

95/96

-40

-20

0

96/97

-40

-20

0

97/98

-40

-20

0

98/99

-40

-20

0

99/00

-40

-20

0

00/01

-40

-20

0

01/02

Dec Jan Feb Mar Apr

 SAOZ REPROBUS SLIMCAT

17%

31%

30%

22%

20%

5%

23%

10%

10% Total Ozone Column Reduction (%)

Figure 3-33. Total ozone reduction in the Arctic derived by comparing measurements of the SAOZ network (Ny

Ålesund, Thule, Scoresbysund, Sodankylä, Salekhard, Zhigansk, and Harestua) with passive ozone in REPROBUS

(gray shaded area) and modeled ozone losses (solid line, REPROBUS; dotted line, SLIMCAT). Written percentages on the right side of each panel are the cumulative SAOZ losses. Reproduced from EC (2001).
of ozone. Associated with persistently low temperatures

in the winters of 1994/1995, 1995/1996, 1996/1997, and

1999/2000, large chemically induced ozone reductions of

22 to 31% were observed inside the vortex. The ozone

loss during the warmer winter of 1998/1999 was smaller

and is hardly significant. In the relatively warm winter of

1997/1998, ozone column losses derived from the transport model approach were still significant (20%). In that

winter, temperatures dropped below the PSC threshold

only in very limited geographical regions inside the vortex

and only during short periods (see Figure 3-30). From the

vortex average approach (Figure 3-31) some limited ozone

loss in 1997/1998 is also visible, but only below 450 K.

During that year Match results at 475 K indicate no

significant loss, but results for 450 K and below are not

available from Match for that year.

3.3.2.4 CHEMICAL OZONE LOSS IN THE ARCTIC

WINTER 1999/2000

The winter of 1999/2000 had the largest potential

for PSC formation for at least the last 20 years. Because

of the extensive SOLVE-THESEO 2000 campaign during

the winter 1999/2000, all basic approaches outlined in

Section 3.3.1 were used to study the ozone loss throughout

the winter, resulting in a better characterization of the

ozone losses in 1999/2000 than in any previous winter

and providing a unique opportunity to compare results

from the different techniques.

All approaches identified extensive chemical loss

of ozone. Figure 3-34a shows the evolution of the vortex

averaged ozone loss in a vertical section as determined

with the Match approach. Ozone loss started at altitudes

above 500 K in mid-January. The largest loss rates of

61 ± 4 ppbv per day (vortex average) were observed at

450 K in early March. Rex et al. (2002a) showed that the

vertical structure and the time evolution of the observed

ozone loss agree well with the vertical structure and the

time evolution of observations of high levels of active

chlorine. During the winter of 1999/2000 only approaches

that are based on ozonesonde measurements, like Match

and the vortex average approach, were able to capture the

full extent of the ozone loss, i.e., cover the full altitude

range of the loss and the time period from early January

to late March in vertical resolution. Comparisons of the

results with other approaches are possible for a slightly

shorter period and focus on the region close to 450 K.

Here the results of all approaches agree very well (see

Section 3.3.3).

Figure 3-35 shows the impact of the cumulative

ozone loss through the winter on the vortex-averaged vertical ozone profile at the end of March. In a layer of air

around 18 km altitude the degree of chemical ozone

destruction reached 70%.

Table 3-2 gives an overview of various estimates

for the deficit in the total column amount of ozone due to

chemical loss of ozone. The numbers given are the difference between the actually observed column amount of

ozone and the column amount of ozone that would have

been present at a given day in the absence of chemical

ozone loss, dynamics being equal. For comparable time

periods the agreement between results from all approaches

is within the error bars. The results from the SAOZ/-

transport model study are generally somewhat higher than

the other approaches and have larger uncertainties. The

results of the other approaches agree to better than 20%.

Figure 3-34c shows the evolution of the total column loss

from January through March, as determined by Match.

By the end of March, the chemically induced ozone deficit

amounted to between 90 and 100 DU. This is roughly the

amount of total ozone that has been supplied to the polar

vortex by dynamic effects during the same time, and therefore the total ozone column remained relatively constant

during January to March (Rex et al., 2002a), which is in

contrast to the natural, climatological increase of the

Arctic ozone column during this season.

3.3.3 Consistency Between the Different

Observational Techniques

Comparisons of the different approaches used to

infer Arctic ozone loss are often hampered by the fact that

the altitude range, horizontal extent (vortex definition),

and time periods used in the various published works

are different. These differences are partly unavoidable

because of the constraints of the datasets used. But often

the datasets can be reanalyzed for certain time periods and

regions where they overlap, so that the results can be

directly compared.

For winter 1999/2000 basically all approaches can

be used to calculate the amount of ozone loss that occurred

inside the polar vortex between 20 January and 12 March

in the layer of air that subsided from about 475 to 450 K

during that time. This subsiding layer of air is indicated

in Figure 3-34a by the solid black lines. Figure 3-36 summarizes the accumulated ozone losses as determined for

this layer of air with the various techniques. The average

of the various estimates is 1.65 ppmv of ozone loss for

this specific time period and vertical region. The results

from all techniques are within ±20% of this value; all but

two are within ±10% of the average.

The winter of 1999/2000 was characterized by relatively weak dynamic activity and perhaps less than

average exchange of air across the vortex edge (see

POLAR OZONE

3.50
Section 3.2.1). It is reasonable to assume that for the

winter of 1999/2000 the agreement between different

techniques to estimate the degree of chemical ozone loss

from ozone observations may be better than for dynamically more active winters. However, to investigate this,

Harris et al. (2002) reanalyzed data from past winters

during the 1990s, using different techniques and datasets, focusing on time periods where the datasets overlap.

On the basis of results from many winters, they found an

agreement of generally better than 20% between techniques that use explicit transport calculations. Results

from tracer relation studies showed slightly larger discrepancies compared with these results, when the initial

tracer relation was measured early in fall. This agreement

is 25% or better, when the initial tracer relation was measured in midwinter, e.g., in December.

On the basis of these studies, the results from current estimates of the degree of chemically induced Arctic

ozone losses appear to have an accuracy of about 20%.

3.3.4 The Effect of Denitrification on

Ozone Loss

The effect of denitrification on ozone loss has been

quantified (to some extent) in both hemispheres. In the

Antarctic (complete) denitrification is shown to cause a

10% increase in the column ozone loss (Brasseur et al.,

1997; Portmann et al., 1996).

Evidence for a much more significant impact of

denitrification (see Section 3.2.2.3) on Arctic ozone losses

in recent cold winters has increased. Observational results

indicate that the degree of ozone loss in the Arctic was

significantly amplified by denitrification during the winters of 1994/1995, 1995/1996, and 1999/2000. Using

model studies, Rex et al. (1997) concluded that in winter

1995/1996, observed ongoing chemical ozone loss in certain air masses more than 1 month after the last exposure

to PSCs can only be explained by approximately 80%

denitrification in about half of the air masses inside t half of the air masses inside

the polar vortex. In that winter the heavily denitrified

layer of air was limited to a very narrow vertical region of

less than 1 km thickness at about 20 km altitude.

Waibel et al. (1999) presented measurements of

denitrification in the Arctic winter of 1994/1995, based

on the NOy-versus-N2O relation, and used a chemical

model to conclude that in the denitrified air masses the

degree of ozone loss was enhanced by at least 30% compared with what would have occurred without denitrification. They show that the model results can come close to

the observed ozone loss only when the observed denitrification is taken into account.

POLAR OZONE

3.51

Cumulative ozone loss [ ppmv ]

012

550

500

450

400 Potential temperature [ K ] Ozone column loss [ DU ] 50

100

0 30 60 90

Date [ da of the ear ]

(c)

(a) (b)

0123

Cumulative

ozone loss

[ ppmv ]

550

500

450

400 spring equivalent potential temperature [ K ]

Figure 3-34. (a) Evolution of the cumulative ozone loss in subsiding air masses from Match for 1999/2000.

The subsidence of air is indicated by the black lines (see text). (b) Profile of the cumulative ozone loss on

26 March 2000 from Match. (c) Accumulated chemical ozone loss in the partial column between q = 400 and

575 K from Match. Error bars are 1s. Reproduced from Rex et al. (2002a).
POLAR OZONE

Table 3-2. Comparisons of chemical loss of column ozone, column [O3* – O3] (see Rex et al., 2002a),

inside the Arctic vortex for the winter of 1999/2000 as of the indicated date. N/A indicates that data for

that date are not available.

Data Source Method Reference Chemical Loss of Column Ozone (DU)

5 March 2000 15 March 2000 28 March 2000

OMS balloon Tracer-tracer Salawitch et al. (2002a) 61 ± 14 N/A N/A

(O3 vs. N2O)

HALOE Tracer-tracer Müller et al. (2002) N/A 84 ± 13 N/A

POAM III Vortex-averaged Hoppel et al. (2002) 51 ± 11 67 ± 11 N/A

satellite descent

SAOZ Transport model EC (2001) 85 ± 24 98 ± 25 101 ± 30

network

Ozonesondes Match Rex et al. (2002a) 53 ± 11 71 ± 12 88 ± 13

3.52

10

15

20

25

02468 10

0

5

30

Altitude [ km ]

Ozone concentration [ 1012 cm-3 ]

(a)

(b)

Ozone concentration [ 1012 cm-3 ]

Typical Antarctic

Ozone Loss

0 2 4 6 8 10

Ozone concentration [ 1012 cm-3 ]

Arctic Ozone Loss in

Winter 1999/2000

Altitude [ km ]

10

15

20

25

0

5

30

Altitude [ km ]

Figure 3-35. Comparison of ozone loss in the Arctic winter of

1999/2000 and typical Antarctic ozone loss. (a) Concentration

profile of O3 during late winter for observations from ozonesondes

between 20 and 30 March 2000 in the Arctic (red). The profile of

O3* (blue), the abundance of O3 expected in the absence of any

chemical loss, was estimated by allowing the early-winter vortex

average O3 profile from ozonesondes to descend by amounts

based on cooling rates from the SLIMCAT model. For this calculation, the mixing ratio of O3 was assumed to be conserved during

descent and was converted to concentration in the last step. The

profiles are shown for the altitude and pressure the air would have

been at on 25 March. The difference between the profile of O3*

and O3 is hatched. This is an estimate, based on the vortex

average approach, for the accumulated chemical loss of ozone

during the winter. The black lines represent a typical late-March

ozone profile above and below the vertical region where significant ozone loss occurred. Adapted from Rex et al. (2002a). (b)

Illustration of the typical degree of ozone loss in the Antarctic.

Typical midwinter (15 July 1997, blue) and late winter (13 October

1997, red) ozone profiles measured inside the Antarctic polar

vortex at the Neumayer station at 71°S are shown. This panel is

for illustration only, and no attempt has been made to correct for

transport effects, which are weaker in the Antarctic compared with

the Arctic. Poleward and downward transport between July and

October supplies ozone to the Antarctic vortex, and therefore the

hatched area in panel (b) is a conservative estimate of the chemical loss.
Gao et al. (2001b) present measurements of varying

degrees of denitrification (also based on the NOy-versusN2O relation) in different areas of the polar vortex, as

characterized by fixed ranges of N2O and potential

temperature. Succeeding observations of the rate of ozone

loss indicate largest loss rates at the N2O levels that were

most severely denitrified (about 30% larger losses than in

less denitrified air). Gao et al. (2001b) show that these

differences in the ozone loss rate cannot be explained by

differences in solar exposure or initial chlorine activation

and, hence, are most likely a result of the denitrification.

These observational studies have shown that denitrification in cold Arctic winters can cause up to a 30%

increase in ozone loss at a given altitude, a result that is

confirmed by model studies (Chipperfield and Pyle, 1998;

Tabazadeh et al., 2000; Drdla and Schoeberl, 2002). The

more pronounced effect of denitrification on ozone loss

in the Arctic, compared with the Antarctic, is the result of

higher temperatures in the Arctic. In the Antarctic, reactivation of chlorine out of the reforming ClNO3 reservoir

via reactions on cold liquid aerosol particles (e.g., ClNO3 +

H2O) can sustain a high level of active chlorine in spring

(Portmann et al., 1996). This is much less effective in the

Arctic, since the heterogeneous reaction of ClNO3 + H2O

is much slower at the higher temperatures typical for the

Arctic spring. Therefore, the lifetime of active chlorine is

strongly dependent on the rate of formation of ClNO3 and,

hence, the abundance of HNO3. The overall effect of denitrification on Arctic column ozone loss depends on the

vertical range of severe denitrification. Currently, the

effect of denitrification on Arctic ozone is limited to the

altitude range of ~18 to 21 km, where most parcels are

POLAR OZONE

3.53

0 1 2

Ozone loss at e one loss at eΘ=450 K between 20 J =450 K between 20 Jan and 12 Mar 2000 (ppmv) an and 12 Mar 2000 (ppmv)

1.65 ppmv ± 20%

Ozone vs. PV/Θ- LARC DIAL - LARC DIAL

Grant et al. (2002) ant et al. (2002)

Vortex average descent - PO rage descent - POAM III

Hoppel et al. (2002) Hoppel et al. (2002)

Ozone vs. PV/Θ - ozonesondes onesondes

Lait et al. (2002) Lait et al. (2002)

Müller et al. (2002) ler et al. (2002)

Ozone vs. chemical tracer HALOE and OMS . chemical tracer HALOE and OMS

Rex et al. (2002) x et al. (2002)

Match - oz Match - ozonesondes onesondes

Richard et al. (2002) Richard et al. (2002)

Ozone vs. chemical tracer ER-2 . chemical tracer ER-2

Salawitch et al. (2002) witch et al. (2002)

Ozone vs. chemical tracer - OMS and ER-2 . chemical tracer - OMS and ER-2

Schoeber Schoeberl et al. (2002) l et al. (2002)

Vortex average descent - oz rage descent - ozonesondes onesondes

Schoeber Schoeberl et al. (2002) l et al. (2002)

Vortex average descent - PO rage descent - POAM III

Schoeber Schoeberl et al. (2002) l et al. (2002)

Trajector ajectory ensemble advection - oz ection - ozonesondes onesondes

Sinnhuber et al. (2001) Sinnhuber et al. (2001)

passive SLIMCA e SLIMCAT run - Ny un - Ny Ålesund oz lesund ozonesondes onesondes

Nov 1999 -11/14 Mar 2000 v 1999 -11/14 Mar 2000

ozone/tracer studies ozone/tracer studies

individual transport individual transport

calculations calculations

bulk transport bulk transport

calculations calculations

Figure 3-36. Overview of the results of various approaches to estimate the amount of Arctic ozone loss

(ppmv) between 20 January and 12 March 2000. The results are stated for the layer of air that subsided from

about 475 K at the beginning of the period to 450 K at the end. Results from studies that use the relation

between ozone and an inert tracer are plotted in blue; studies that rely on bulk transport calculations are

plotted in red; and results from the approach that relies on individual transport calculations are plotted in

green. The error bars represent the 2-standard-deviation uncertainty of the loss estimate as estimated by the

individual investigators. The average of the various estimates is 1.65 ppmv ± 20%. (OMS = Observations of

the Middle Stratosphere; LARC DIAL = NASA Langley Research Center Differential Absorption Lidar; other

acronyms defined in text.) Based on Newman et al. (2002).
shown to be severely denitrified in cold winters (Hintsa

et al., 1998; Kondo et al., 2000; Fahey et al., 2001).

Microphysical studies have shown that the vertical range

of severe denitrification in the Arctic lower stratosphere

is very sensitive to the temperature (Waibel et al., 1999;

Tabazadeh et al., 2001).

It is important to note that current denitrification

schemes (e.g., in 3-D CTMs) have difficulties correctly

representing the degree of denitrification in cold Arctic

winters (e.g., Davies et al., 2002; see also Section 3.2.2.3).

In light of the recent results showing that extensive denitrification (up to 80%) occurred in cold Arctic winters and

that it had significant impact on the degree of ozone loss

in those years, the correct representation of denitrification

remains one of the major challenges for 3-D CTMs, when

they are used to study the variability of chemical ozone

loss in the Arctic (e.g., Chipperfield and Jones, 1999) or

to predict future ozone losses in a potentially changing

climate.

Dehydration, unlike denitrification, can moderate

ozone loss for two reasons (Portmann et al., 1996;

Chipperfield and Pyle, 1998). First, in a drier atmosphere

it is harder for PSCs to form. Second, heterogeneous reaction rates that lead to active chlorine production drop

exponentially with decrease in relative humidity.

Sensitivity studies show that dehydration (to the level of

ice saturation) in the Antarctic can decrease column ozone

loss by about 20% (Portmann et al., 1996; Brasseur et al.,

1997). No large-scale model calculations have yet been

performed to evaluate the role that dehydration may play

in Arctic ozone loss and recovery in the future. However,

it is unlikely that climate change in the near future could

cause extensive dehydration in the Arctic region. Some

air mass trajectory statistical analyses indicate that even a

substantial cooling of lower stratospheric temperatures

(by 3 to 4 K) is still insufficient to trigger the occurrence

of severe dehydration in the Arctic vortex (Tabazadeh et

al., 2000).

3.3.5 Model Studies of Arctic Ozone Loss

Model investigation of polar ozone loss was the

subject of intense research in recent years. Most studies

concentrated on the Arctic region because of the high

interannual variability of the Arctic ozone loss in relation

to the year-to-year meteorological conditions. Threedimensional CTMs, which proved to be particularly well

fitted to the non-zonal character of the Arctic polar vortex,

were used to estimate the overall degree of polar ozone

loss for several winters. For more specific studies or

highly constrained comparisons between models and

ozone loss observations, photochemical box models were

used.

3.3.5.1 CHEMICAL TRANSPORT MODELS

Studies using the transport model approach to estimate ozone loss from ozone measurements typically

include a comparison with the ozone loss calculated by

the chemistry module of the model. Chemical ozone loss

inferred from the POAM II and III measurements was

compared with that obtained from the REPROBUS model

(Deniel et al., 1998, 2000). In the same way, total ozone

measurements by the SAOZ network were compared with

REPROBUS and SLIMCAT simulations for various

Arctic winters from 1993/1994 (Figure 3-33; Goutail et

al., 1999; EC, 2001). The agreement between the

observed ozone loss and the model result varies from

winter to winter. Two points can be seen in Figure 3-33.

First, the overall interannual variability of the Arctic ozone

loss is reasonably well represented by the models. The

models reproduce a large fraction (about 60 to 100%) of

the overall Arctic ozone loss. Second, in winters when

substantial ozone loss was observed during January

(1994/1995, 1995/1996, and 1999/2000), the models fall

short of reproducing this January loss. Typically, by the

end of January, only about 50% of the observed loss is

accounted for by the models. The loss in January is partly

why the overall loss at the end of the winter is sometimes

underestimated (e.g., 1994/1995). In other winters the

January loss contributes only a minor fraction to the

overall loss (e.g., 1999/2000) or the model overestimates

the loss rate later during the winter, and therefore the

overall loss at the end of the winter is better reproduced

than the time evolution of the loss (e.g., 1995/1996). Also,

in 1995/1996 the simulated vertical distribution of the

ozone loss at the end of March differs from that estimated

from POAM II measurements, with the model resulting

in larger losses at lower altitudes and smaller losses at

higher altitudes compared with the observations (Deniel

et al., 2000). It appears that a good agreement between

observations and modeled total ozone loss at the end of

the winter alone may be fortuitous and does not necessarily prove that the ozone loss mechanisms in the model

are well reproduced.

Extensive modeling studies were performed as part

of SOLVE/THESEO 2000 to estimate the ozone loss in

the winter of 1999/2000. Sinnhuber et al. (2000) compared the chemical ozone loss estimated with SLIMCAT

with that derived from the model ozone passive tracer and

ozonesonde observations at Ny Ålesund (Figure 3-37).

They found good agreement between the modeled ozone

and observations, both indicating more than 2.5 ppmv

ozone destruction by late March, corresponding to 70%

ozone loss at the 450 K isentropic level, the largest ozone

loss ever produced by SLIMCAT. The reason for the large

loss of ozone in the model was extensive formation of

POLAR OZONE

3.54
large denitrifying ice particles by the model’s microphysical scheme. But the large-scale formation of ice clouds

in the model was the result of a cold bias in the UKMO

temperature fields used in SLIMCAT and is not consistent with observations during SOLVE/THESEO 2000. A

later study performed by Davies et al. (2002) focused on

the effect of denitrification on Arctic ozone depletion

during the same winter using SLIMCAT and a denitrification scheme based on the formation of large NAT particles (see Section 3.2.2.3). When the model was forced by

UKMO analyses, the ozone losses derived inside the

vortex were very close to those derived by Sinnhuber et

al. When ECMWF temperatures were used that were

closer to observations in winter 1999/2000, the modeled

active chlorine was in better agreement with ER-2 measurements, but the derived ozone loss was smaller than in

observations and in Sinnhuber et al. (2002).

The KASIMA (Karlsruhe SImulation model of the

Middle Atmosphere) CTM driven by ECMWF analyses

was compared with ozone measurements by a Fourier

transform infrared (FTIR) spectrometer and a millimeter

wave radiometer in Kiruna for winter 1999/2000. The

modeled total ozone loss underestimates the observations

by 30% and 20%, respectively (Kopp et al., 2002).

The Langley Research Center (LaRC) Lagrangian

chemical transport model (LCTM) was used in conjunction with HALOE and POAM III satellite observations to

simulate the large-scale photochemical evolution of the

Arctic vortex in 1999/2000 from a vortex ensemble of air

mass trajectories using UKMO analyses (Pierce et al.,

2002; see also Appendix 3Afor satellite data descriptions).

The model shows significant denitrification within the

vortex in late December and early January. A significant

overprediction of the level of chlorine activation is found

in early March, but the predicted peak ozone loss rate is

in good agreement with that inferred from the Match

campaign during the same period. Conversely, it can be

concluded that for a realistic level of active chlorine, the

model would have significantly underestimated the

observed ozone loss rate during that period.

Grooß et al. (2002) report simulations with the

Chemical Lagrangian Model of the Stratosphere

(CLaMS). This model simulates the dynamics and chemistry of multiple air parcels along their trajectories that

are determined from ECMWF winds. The model includes

mixing between neighboring air parcels. In this model

study, the degree of denitrification was described from

observations by using observed relations between NOy

and N2O and the temperature history based on ER-2 measurements. The simulation was initialized on 10 February,

and the ozone loss during the mid-February to mid-March

period (up to 60% at 425 to 450 K) agrees roughly with

estimates from observations.

The various 3-D model studies focusing on winter

1999/2000 reveal a consistent picture. The ozone loss

after mid-February is well reproduced if the degree of

denitrification in the model is correct, whether by coincidence as in Sinnhuber et al. (2000) or because it was

specified from observations as in Grooß et al. (2002).

Currently, 3-D CTMs are not able to reliably reproduce

observed denitrification. This deficit limits the current

ability to reliably reproduce the degree of ozone loss in

cold Arctic winters.

3.3.5.2 SPECIFIC MODEL STUDIES

Several modeling studies using box models were

conducted to compare specific ozone loss observations

with model results and test our understanding of the chemical processes involved in the loss. These calculations are

performed specifically for the air masses in which the

ozone losses have been observed. Hence, the temperature and solar zenith angle history in these studies are

much more constrained than in comparisons of vortexaveraged ozone losses. Becker et al. (1998, 2000) performed box model simulations along the trajectories of

the Match dataset for 1991/1992 and 1994/1995. This

analysis is illustrated in Figure 3-38, which compares the

POLAR OZONE

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60 30 0 30 60 90 120

1

2

3

4

Date [day of year 2000] Ozone [ppm]

Ny Ålesund

450K

NOV DEC JAN FEB MAR APR

SLIMCAT 

passive ozone

ozone sondes 

Figure 3-37. Ozonesonde measurements at Ny Ålesund (79°N) compared with the SLIMCAT threedimensional transport model output for Ny Ålesund

at the 450-K isentropic level. Ozone loss is indicated

by the difference between the SLIMCAT’s passive

tracer (gray line) and SLIMCAT’s ozone output (black

line). Adapted from Sinnhuber et al. (2000).
ozone loss rates derived from simulations along the Match

trajectories with the corresponding Match results at four

potential temperature levels between 440 and 490 K for

the winter of 1994/1995. It shows that the model underestimates the ozone losses observed by Match in late

January by up to a factor of 2 above 475 K. During the

other months, the observed losses are also underestimated

by the model but within the large uncertainties of the

model, mainly linked to the extent of denitrification.

Match ozone loss rates were also compared with

SLIMCAT simulations for the winters of 1994/1995 and

1995/1996 (Kilbane-Dawe et al., 2001). The study suggests that Match may have overestimated the ozone loss

rates above 525 K in January 1995 because of deficiencies in the ECMWF wind fields close to the top level of

the ECMWF assimilation model (the top level has shifted

to higher levels since then). In January 1995 at levels

below 525 K and in January 1995/1996, SLIMCAT generally underestimated Match ozone loss rates by about 30

to 50%. It was found that the SLIMCAT photochemistry

POLAR OZONE

3.56

Day of the Year 1995

0

4

8

0

4

8

 Ozone Loss Rate [ppb/Sunlit Hour] 0

4

8

0

4

8

12 41 86 63 18 36 40 49 51 39

54 92 65 25 46 47 51 52 37 45 27

66 89 61 26 50 52 61 56 31 38 24

48 83 66 26 39 45 59 43 14 24 15

Day of the Year 1995

 20

0

20

40

60

80

0

20

40

60

80

Ozone Loss Rate [ppb/Day] 0

20

40

60

80

0

20

40

60

80 19 41 86 63 18 36 40 49 51 39

25 54 92 65 25 46 47 51 52 37 45 27

40 66 89 61 26 50 52 61 56 31 38 24

26 48 83 66 26 39 45 59 43 14 24 15

Θ = 440 K

Θ = 470 K

Θ = 480 K

Θ = 490 K

Θ = 440 K

Θ = 470 K

Θ = 480 K

Θ = 490 K

0 30 60 90 0 30 60 90

Figure 3-38. Time dependence of the ozone loss rates per sunlit hour (left panels) and the ozone loss rates

per day (right panels) at various potential temperature levels in the 1994/1995 Arctic winter. Results of the

Match analysis (open black circles) are compared with model results obtained with denitrification (solid gray

circles) or without denitrification (solid black circles). Each point was deduced from Match trajectories with

less than a ±7-day time interval and a ±10-K potential temperature interval. The number of matches contributing to each point is specified at the upper edge of each panel. The error bars correspond to the 1σ uncertainty of the linear regression.
was least able to reproduce observed ozone losses when

low temperatures coincide with high solar zenith angles.

Woyke et al. (1999) used the tracer relation

approach to quantify ozone loss in air masses that had

been probed by a balloon payload providing observations

of ClO, BrO, ozone, and long-lived tracers, on 3 February

1995. They used box model runs constrained by ClO and

BrO concentrations observed by the balloon, to calculate

the ozone loss throughout January along the back trajectories of the air masses. Using this highly constrained

approach, they could explain only half of the observed

ozone loss.

These results confirm that ozone losses observed

during cold Arctic Januarys are currently not understood.

3.3.6 Quantifying Antarctic Ozone Loss

As emphasized in Section 3.1, Antarctic ozone

depletion has been monitored by ground-based and satellite measurements since the mid-1980s. However, relatively few recent studies have concentrated on a detailed

quantification of Antarctic ozone loss rates with stateof-the-art approaches to separate chemical loss from

dynamical impacts. Hence, our quantitative knowledge

of Antarctic ozone loss rates has not been as extensively

tested as in the Arctic. The quantification of the accumulated overall ozone loss in the Antarctic is not challenging,

because by the end of the winter, ozone is basically completely lost in a broad vertical region.

Hofmann et al. (1997) quantified Antarctic ozone

loss from the analysis of 10 years of ozonesonde measurements at the South Pole and made recommendations

for the detection of the recovery of the Antarctic ozone.

Indicators for recovery include an end to springtime ozone

depletion at 22 to 24 km and a 12- to 20-km ozone partial

column value of more than 70 DU on 15 September.

Bevilacqua et al. (1997) used POAM II ozone observations above Antarctica to derive vortex average ozone loss

rates in August and September from 1994 to 1996 in the

450- to 800-K potential temperature range. Significant

loss of ozone was found over the whole three-year period,

except in 1994, when ozone loss was not observed in

August, due to the sampling of the POAM instrument.

The largest loss rates were found in September 1996 where

they reached 0.1 ppmv/d below 500 K. From the analysis

of the temporal evolution of total ozone at the Faraday

station together with model calculations, Roscoe et al.

(1997) showed that the ozone chemical depletion starts in

June at the sunlit vortex edge, and Waters et al. (1999)

showed that enhanced abundances of ClO are observed

on the sunlit edge of the Antarctic vortex by late May or

early June. Ozone loss rates were evaluated above the

Antarctic station of Dumont d’Urville (66.4°S, 140°E)

from ozonesonde and lidar measurements on an interannual basis (Godin et al., 2001). Interpretation of the data

required careful analysis of PV-equivalent latitude to

determine whether each observation was inside, in the

edge or outside the vortex at different isentropic levels.

Measurements inside the vortex showed complete ozone

destruction from 400 to 500 K, with ozone loss rates

reaching 0.06 ppmv/d in the late-August to September

period.

3.3.7 Model Studies of Antarctic Ozone

Loss

The few recent model studies of the Antarctic

ozone loss generally point to an agreement between

models and observations. The SLIMCAT model was used

to study the austral stratosphere in winter and spring 1996

together with ozonesonde measurements from various

Antarctic stations (Lee et al., 2000). The model shows

very good agreement with measured ozone values, and

both the model and observations show that chemical

ozone depletion follows the edge of polar night with little

mixing poleward until the terminator reaches 80°S. In a

follow-up study, Lee et al. (2001) analyzed the isentropic

transport processes within the Antarctic polar vortex.

Their calculations indicate two distinct regions within the

vortex: a strongly mixed vortex core and a broad ring of

weakly mixed air that remains isolated from the core

between late winter and midspring and where the ozone

loss is not complete. This result has an implication for

the recovery of Antarctic ozone, since a cooling of the

stratosphere could enhance the ozone loss in the edge

region and delay the ozone recovery. In another study of

Antarctic ozone loss, Wu and Dessler (2001) tested the

current understanding of polar ozone chemistry with MLS

version 4 measurements of ozone and ClO. By comparing

the observed ozone loss estimated from the MLS ozone

evolution at 465 K with a modeled ozone loss inferred

from the simultaneous ClO measurements and a fixed BrO

mixing ratio, they find a good agreement between both

methods. However, MLS version 5 data, which has a

better vertical definition of the ClO profile, resulted in a

significant reduction of the ClO concentrations at 465 K

compared with the version 4 data that had been used in

Wu and Dessler (2001). This reduction would lead to a

reduction in the modeled ozone loss rate in Wu and

Dessler (2001). A slight change in the ozone profile in

MLS version 5 data compared with version 4 data would

also reduce the ozone loss rate deduced from observations. A quantitative study would be required to assess

how the conclusions of Wu and Dessler (2001) would

change if MLS version 5 data had been used (see

Appendix 3A for MLS data description).

POLAR OZONE

3.57
3.3.8 Conclusion

Arctic chemical ozone losses during the last decade

have been determined by a variety of approaches, and

ozone loss rates are now better quantified in the Arctic

than in the Antarctic. The uncertainty of state-of-the-art

approaches to quantify Arctic ozone losses from ozone

observations is below 20% for local losses between 400

and 550 K potential temperature and perhaps somewhat

larger for total column loss estimates. Large interannual

variability of the Arctic ozone loss, ranging from 0 to 70%

loss at about 20 km for individual winters during the past

decade, is driven by the variable extent of temperatures

low enough for PSC formation in a given winter. Global

CTMs reproduce a large fraction (60 to 100%, depending

on the winter) of the observed ozone loss in the Arctic and

its variability. The largest uncertainties are due to the current unrealistic representation of denitrification processes

in 3-D CTMs and unexplained ozone losses during cold

Arctic Januarys. These uncertainties currently prevent

reliable predictions of future Arctic ozone losses in a

potentially changing climate.

3.4 CAUSES OF POLAR STRATOSPHERIC

TEMPERATURE TRENDS

3.4.1 Introduction

In the previous Assessment (WMO, 1999), it was

recognized that the future development of the ozone layer

does not depend just on changes in stratospheric halogen

loading but also, very importantly, on a number of other

factors connecting chemistry and climate. The observations of temperatures are discussed in Section 3.1.2, and

the causes of trends in polar stratospheric temperatures

are discussed here: the role of changes in well-mixed

greenhouse gases (WMGHGs), ozone, water vapor, and

aerosol particles are reviewed. Solar effects are also

noted. The onset of low temperatures during the polar

winter/spring, their duration, interannual variations, and

the statistical significance of trends, over the past two

decades, are issues that impact upon our knowledge of the

chemistry-climate interactions and the detection and attribution of climate change in the polar stratosphere due to

ozone and other greenhouse gases.

Observations, from radiosondes and satellites,

have shown a general cooling of the polar lower stratosphere over the last few decades (WMO, 1999; Ramaswamy et al., 2001). For a number of reasons detection

and attribution of temperature change in the lower stratosphere may be easier than at the surface (IPCC, 1996).

First, the observed temperature change in the stratosphere is large and the response time of the stratosphere

is shorter, compared with the surface. There are relatively good satellite observations of both temperature and

the important atmospheric constituents over the last few

decades, corresponding to the timing of polar ozone

depletion. There are also potentially fewer relatively

uncertain mechanisms involved in stratospheric temperature change; many of the large and uncertain surface

radiative forcings, such as the anthropogenic sulfate

aerosol forcing, are expected to have a minimal effect on

stratospheric temperatures. Further, the magnitude of the

response in the stratosphere to a given mechanism has

been shown to be reasonably well approximated by purely

radiative processes, and therefore may be better quantifiable than the surface temperature response (Ramaswamy

et al., 2001). However, the WMO (1999) ozone assessment acknowledged that the large variability of temperatures, particularly in the Arctic winter and spring, and a

possible stratosphere-wide trend in stratospheric water

vapor complicate the attribution issue. It is now recognized that ozone and WMGHG changes cannot be considered in isolation, and there is an increasing acknowledgement that it is important to attempt to quantify the

feedbacks between temperature change, chemistry, and

stratospheric dynamics, to understand better the stratospheric temperature response.

Since the previous Assessment there has been

improved quantification of atmospheric constituent

changes and development of more sophisticated stratospheric models, especially coupled chemistry-climate

general circulation models. These have provided important insights into our understanding of polar temperature

changes in the lower stratosphere.

This section uses the updated lower stratospheric

high-latitude temperature trends (discussed in Section

3.1.2) and discusses the recent modeling efforts that have

attempted to understand them. It concentrates on the

analysis and understanding of past decadal-time-scale

trends in the polar lower stratosphere; the upper stratospheric response is often more independent of latitude and

is discussed in the global ozone chapter (Chapter 4).

3.4.2 Modeling Techniques

Several different types of model have been adopted

for the study of stratospheric temperature change. Two

types of commonly used models are briefly assessed here.

Figure 3-39 compares the range of model results to observations, and Figure 3-40 shows individual model temperature trends.

POLAR OZONE

3.58
3.4.2.1 FIXED DYNAMICAL HEATING (FDH) MODELS

Fixed dynamical heating (FDH) models and seasonally evolving fixed dynamical heating (SEFDH)

models (WMO, 1999; IPCC, 2001) employ a method of

calculating temperature changes using only a radiative

transfer model. They have been shown to agree well with

calculations using GCMs (Rosier and Shine, 2000;

Ramaswamy et al., 2001), and, compared with these, they

are generally faster and allow the use of more sophisticated radiative transfer schemes. In contrast to the FDH

technique, SEFDH techniques include a calculation for

the time evolution of temperature and have been shown

to improve the temperature response in the high-latitude

polar stratosphere, compared with GCM integrations

(Forster et al., 1997; Rosier and Shine, 2000). This was a

region where the equilibrium temperature response calculated with FDH models overestimated the cooling

resulting from short-term polar ozone depletion. Both

FDH and SEFDH techniques are unable to model the

response of atmospheric dynamics.

3.4.2.2 GENERAL CIRCULATION MODELS (GCMS)

The GCM-Reality Intercomparison Project for

SPARC (GRIPS) is currently performing an assessment

of the performance of current middle-atmosphere GCMs

(Pawson et al., 2000). Preliminary analysis of their results

suggests that all models have a cold bias, at most levels in

the troposphere and stratosphere, which may be indicative of errors in the radiative transfer, or input data

(Pawson et al., 2000). This is particularly pronounced at

the poles; one of the largest uncertainties remains the parameterization of small-scale gravity waves (see Sections

3.2.1.2 and 3.5.2.1). The upper boundary and vertical resolution may also be important for an accurate simulation

(see Section 3.5.2.3). Because of large interannual variability, particularly in the Arctic temperatures (see

Sections 3.1.2 and 3.5.2), either many transient integrations or many years of equilibrium experiments need to

be performed for statistically significant trend calculations. It must be noted, however, that the real atmosphere

has evolved through only one specific realization and that

the global observational record spans only ~20 years.

Experiments have been performed with both prescribed

changes in atmospheric constituents and coupled

chemistry-climate GCMs (discussed in more detail in

Section 3.5). Although coupled chemistry-climate GCMs

allow interaction between radiation, chemistry, and

dynamics, they have two main drawbacks when

attempting to attribute stratospheric temperature change

to a particular cause. First, their complexity makes them

computationally expensive; it is therefore difficult to run

them for long enough periods to produce statistically significant trends. Second, their simulation of the ozone

change is imperfect; this is to be contrasted with GCMs

that employ the ozone trends inferred from observations

to determine the temperature response.

3.4.3 Causes of Polar Stratospheric

Temperature Trends

This section assesses the role of different mechanisms on polar lower stratospheric temperature trends. A

variety of recent model results are compared with the

observed temperature trends (discussed in Section 3.1.2;

see also Figure 3-9). Figures 3-39 and 3-40 summarize

the findings for the lower stratospheric temperature

response at 70°N and 70°S to (1) stratospheric ozone

changes, (2) WMGHG changes, (3) combined changes in

WMGHGs and stratospheric ozone changes, and (4) stratospheric water vapor changes. Annual and seasonal temperature trends are shown at pressures of 50 hPa and 100

hPa. Results are taken from studies by Ramaswamy and

Schwarzkopf (2002), Rosier and Shine (2000; updated by

M. Bourqui), Langematz et al. (2002), Smith (2001),

Butchart et al. (2000), Shindell (2001), Austin (2002),

Schnadt et al. (2002), and Forster and Shine (2002).

Several factors affect the interpretation of the comparisons of the observations and model simulations presented

in this section:

(1) The satellite measurements comprise radiances from

a range of altitudes. This introduces some uncertainty

when model results at 50 and 100 hPa are compared

with the satellite observations.

(2) This section examines a single latitude belt in the

polar regions where there is a large dynamical variability (see Figure 3-8).

(3) The polar trends, especially for the winter/spring

season, are influenced by the end-year considered for

the analysis.

3.4.3.1 OZONE

The stratospheric cooling over Antarctica has been

shown previously to be very well correlated with ozone

losses (WMO, 1999; Randel and Wu, 1999a). Shine

(1986) was the first to show that such a stratospheric

cooling could be due to the direct radiative response of

the ozone loss. GCM and radiative model studies since

then have largely confirmed this early work (Mahlman et

al., 1994; Ramaswamy et al., 1996; WMO, 1999). In the

NH, studies have also found a correlation between

POLAR OZONE

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POLAR OZONE

3.60

50 hPa, 70°N, MAM

-4

-3.5

-3

-2.5

-2

-1.5

-1

-0.5 Satellite Berlin RAOB Russia CPCO3 WMGHG+O3 ALL K/decade 1979 -1998 100 hPa, 70°N, MAM

-4

-3.5

-3

-2.5

-2

-1.5

-1

-0.5 K/decade

50 hPa, 70°S, SON -4

-3.5

-3

-2.5

-2

-1.5

-1

-0.5 Satellite RAOB CPCO3 WMGHG+O3 ALL K/decade

100 hPa, 70°S, SON -4

-3.5

-3

-2.5

-2

-1.5

-1

-0.5 K/decade 1979-1998

50 hPa, 70°N, Annual

-4

-3.5

-3

-2.5

-2

-1.5

-1

-0.5 K/decade

50 hPa, 70°S, Annual -4

-3.5

-3

-2.5

-2

-1.5

-1

-0.5 K/decade

(-4.68)

100 hPa, 70°N, Annual

-4

-3.5

-3

-2.5

-2

-1.5

-1

-0.5 K/decade

100 hPa, 70°S, Annual

-4

-3.5

-3

-2.5

-2

-1.5

-1

-0.5 K/decade 1979-1998 1979-1998

(-5.12)

Figure 3-39. Comparison of observed and modeled temperature trends at 70°N and 70°S for the spring seasons and annual averages. Trends are presented at 50 hPa and 100 hPa. Observed trends are represented

by blue bars and discussed in Section 3.1.2. All observed trends are significant at, at least, the 1s level. The

average modeled trends from Figure 3-40 are shown as the red bars with the range of model results illustrated

by the black “uncertainty bars.” Model trends are shown for ozone changes alone, WMGHG and ozone

changes, and “All,” where model estimates of the effects of stratospheric water vapor have been added to the

WMGHG and ozone model trends. All observed trends are for 1979-2000 unless otherwise indicated (see

Ramaswamy et al. (2001) for details regarding the observational datasets).
POLAR OZONE

3.61

MSU +SSU Ramaswamy Rosier Langematz Smith

A. STRATOSPHERIC OZONE

MSU +SSU Ramaswamy Rosier

Langematz Butchart Shindell

MSU +SSU Ramaswamy Rosier Langematz

Austin Shindell Schnadt

MSU +SSU Forster Shindell Smith Smith 2D

B. WMGHG

C. WMGHG AND STRATOSPHERIC OZONE

D. STRATOSPHERIC WATER VAPOR

70°N

70°N

70°N

70°N

70°S

70°S

70°S

70°S

DJF MAM JJA SONAnnual DJF MAM JJA SONAnnual

DJF MAM JJA SONAnnual DJF MAM JJA SONAnnual

DJF MAM JJA SONAnnual DJF MAM JJA SONAnnual

DJF MAM JJA SONAnnual DJF MAM JJA SONAnnual

-4

-3

-2

-1

0

1

2

-3

-2

-1

0

1

2

-4

-3

-2

-1

0

1

2

-3

-2

-1

0

1

Temperature Trend (K/decade)

Figure 3-40. The seasonal and annual combined 50- and 100-hPa temperature trends at 70°N and 70°S,

corresponding to one-third of the 50-hPa trend added to two-thirds of the 100-hPa trend, using the modelderived trends discussed in this section. The modeled temperature trends are for (a) stratospheric ozone

changes; (b) WMGHG changes; (c) WMGHG and stratospheric ozone changes; and (d) stratospheric water

vapor changes. On each panel, 2 error bars are plotted for the observed satellite trends, and the horizontal

dotted lines indicate the zero trend. References for models are given in Section 3.4.3.
springtime Arctic temperatures and ozone (WMO, 1999;

Randel and Wu, 1999a). However, the correlations are

not as strong, and the cooling observed during the Arctic

winter is not expected from a simple radiative response to

the ozone loss (Randel and Wu, 1999a). In addition,

reductions of planetary wave driving reduce the strength

of the residual circulation, leading to a cooling trend. This

also weakens the transport of ozone-rich air to the polar

lower stratosphere, and could give larger heterogeneous

loss rates. Hence, a correlation of temperature and ozone

does not necessarily imply a causal linkage of ozone loss

with temperature.

Since the previous Assessment several studies have

employed the monthly averaged vertical trend data, based

on ozone observations over the period 1979 to 1997

(SPARC, 1998), with most studies employing the combined trend dataset of Randel and Wu (1999b). The results

of these prescribed ozone change studies are shown Figure

3-40a for the NH (left panel) and SH (right panel). The

calculations presented in the figure had the benefit of more

detailed ozone-trend vertical resolution compared with

previous work. The ozone dataset employs stratosphericonly trends derived from the Syowa (69°S) and Resolute

(75°N) ozonesondes as representative of the polar regions.

Figure 3-41 shows an illustration of the annual cycle of

model-derived 100-hPa trends from the Berlin model and

compares them with trends derived from reanalysis data.

It is noted that trends from reanalysis data (Figure 3-41,

panel c) must be interpreted with caution (Chapter 5 in

WMO, 1999). The Antarctic and Arctic response are discussed separately below.

Antarctic

All model studies report an ozone-induced cooling

at 70°S for most seasons, with the largest cooling in all

models occurring in September, October, and November

(SON) and December, January, and February (DJF) (right

panel of Figure 3-40a). Langematz et al. (2002) and

Ramaswamy and Schwarzkopf (2002) find that the largest

cooling at 100 hPa occurs in DJF not SON (see Section

3.4.4). There is a dynamically induced heating in the

middle atmosphere, which adds to the radiative heating

owing to more upwelling longwave radiation initiated by

the depletion of ozone in the lower stratosphere. This

effect is simulated by several models (Kiehl et al., 1988;

Mahlman et al., 1994; Ramaswamy et al., 1996; Rosier

and Shine, 2000) and is consistent with observations

(Ramaswamy et al., 2001), but it is not statistically significant, in either models or observations.

Compared with the observed trends, the Langematz

et al. (2002) response underestimates the SON cooling at

50 hPa, but gives a better fit with the MSU-4 observations

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Jul Aug Sep Oct Nov Dec Jan Feb Mar Apr May Jun Jul

-90

-60

-30

Eq.

30

60

90

Jul Aug Sep Oct Nov Dec Jan Feb Mar Apr May Jun Jul

Jul Aug Sep Oct Nov Dec Jan Feb Mar Apr May Jun Jul

-90

-60

-30

Eq.

30

60

90

-90

-60

-30

Eq.

30

60

90 Latitude Latitude Latitude

(A)

(B)

(C)

Figure 3-41. Annual cycle of the zonal-mean temperature trend (K/decade) at 100 hPa. (a) Modelderived trends employing the observed 1980 to 2000

ozone trend; (b) same as (a), but with an additional

increase in CO2; and (c) the corresponding 1980 to

2000 linear trend, derived from NCEP reanalysis

data. The contour interval is 1.0 K/decade. Dark

(light) shaded areas denote regions where the trend

is significant at the 99% (95%) confidence level.

Adapted from Langematz et al. (2002).
representative of 100 hPa. The other models tend to overestimate the observed spring and summer cooling at 100

hPa, particularly during DJF (see Section 3.4.4). Caution

needs to be applied to this satellite-model comparison,

because there are uncertainties with regard to the representation of the altitude profile of ozone loss in the models

over the winter/spring period, and with regard to the

interpretation of seasonal trends at 50 and 100 hPa

from satellite data.

Two versions of the same model (Rosier and Shine,

2000, updated; Smith, 2001) produce substantially different trends in SON (Figure 3-40a). Both models are

versions of the Reading Intermediate GCM performing

20-year equilibrium experiments. Differences between

the simulations are also found in the Arctic (see below),

which may imply a difference between the way the ozone

trends are implemented and/or differences in the variability of the model.

Arctic

All four models in Figure 3-40a (left panel) find an

annually averaged cooling in the Arctic. Three of the four

models find the largest cooling in March, April, and May

(MAM), which agrees qualitatively well with the satellite

data. However, this modeled cooling is only about half

the size of the cooling trend in the satellite data. Smith

(2001), in contrast, finds a maximum cooling in DJF and

smaller cooling in MAM, although differences again may

be due to the large variability in the Arctic winter and are

of limited significance. When comparing trends, Graf et

al. (1998), Waugh et al. (1999), Langematz et al. (2002),

and Ramaswamy and Schwarzkopf (2002) make the

important point that the large natural variability in the

Arctic may mean that the modeled and observed trends

are not easily compared (see Figure 3-8).

3.4.3.2 WELL-MIXED GREENHOUSE GASES

(WMGHGS)

In general, considering their radiative effects only,

WMGHG increases are expected to have a much smaller

effect on lower stratospheric temperatures than they do

on temperatures above 30 hPa (Forster and Shine, 1997;

WMO, 1999; Chapter 4 of this Assessment; Ramaswamy

and Schwarzkopf, 2002). This is borne out by the results

presented in Figure 3-40b for the NH (left panel) and SH

(right panel). The Langematz et al. (2002) trend was

derived for CO2 changes only and has been estimated

by taking the difference between the results of their

stratospheric-ozone-and-WMGHG experiment and those

of their stratospheric-ozone-only experiment. The same

method has been adopted in connection with the

Ramaswamy and Schwarzkopf (2002) results. This is in

contrast to the updated Rosier and Shine (2000) study

where WMGHG effects were studied directly. Few of the

modeling studies show significant temperature trends.

Some models are seen to cool their annual-mean lower

stratospheres in response to WMGHG increases. The

magnitude of cooling (~0.3 K/decade) agrees with earlier

FDH calculations (e.g., Ramaswamy et al., 2001). For

further discussion on the role of WMGHGs in annualmean trends see Chapter 4.

3.4.3.3 COMBINED CHANGES IN WMGHGS AND

STRATOSPHERIC OZONE

Since the previous Assessment (WMO, 1999) several experiments with coupled chemistry-climate models

have been performed for different time periods to simulate the combined effect of changes in ozone and

WMGHGs. These experiments are discussed in more

detail in Section 3.5. Here we only examine how the coupled chemistry-climate modeling studies influence the

attribution of high-latitude temperature trends. Results

from four coupled chemistry-climate experiments are presented in Figure 3-40c. Temperature trends are also shown

for two different prescribed WMGHG and ozone experiments and the updated Rosier and Shine (2000) result,

where individual ozone and WMGHG trends have been

added.

Antarctic

The agreement between the different models in

Figure 3-40c (right panel) is encouraging considering (1)

the range of different models employed; (2) the differences between the responses, for models employing the

same ozone trend dataset (Section 3.4.3.1); and (3) the

large uncertainty in trends derived from transient integrations. Because the overall effect of WMGHG changes on

lower stratospheric temperatures is minimal (annual-mean

results in Section 3.4.3.2), there is little difference with

respect to the annual-mean temperature trends from the

stratospheric ozone experiments. Therefore, for all seasons the observed cooling can be explained by either

combined WMGHG and stratospheric ozone changes or

stratospheric ozone changes alone (see Figure 3-39). Most

of the coupled chemistry-climate models and Ramaswamy

and Schwarzkopf (2002) tend to overestimate the

observed 50- to 100-hPa cooling in SON, although most

results are still within the observational uncertainty range.

Most of the models presented in Figure 3-40c (right panel)

were found to simulate both the annual-mean and SON

POLAR OZONE

3.63
change at 50 hPa reasonably well, while all overestimated

the 100-hPa trend (not shown). Austin (2002) found that

the discrepancies in their model were consistent with systematic biases in the vertical profile of the applied ozone

loss. Generally, the models capture Antarctic springtime

cooling, but further improvements are needed to simulate

the magnitude of the cooling and its vertical extent.

Arctic

The models for the Arctic, like those for Antarctica,

are generally consistent with the observed trends, and

inter-model differences are generally smaller than the 2s

uncertainty in the observations (left panel of Figure 3-

40c). At 50 to 100 hPa, all models underestimate the

MAM cooling, which could indicate (1) an underestimate

of the ozone-related cooling; (2) a contribution of a

missing effect, such as increases in stratospheric water

vapor; or (3) an underestimate of the role of natural variability. Some models yield a warming in DJF that is in

contrast to the observations, although this may be statistically insignificant in models. In the annual mean, only

one of the models (a coupled model) is close to the observations; the rest underestimate the observed cooling.

3.4.3.4 STRATOSPHERIC WATER VAPOR

Since the previous Assessment, a major development has been the burgeoning interest in the role of stratospheric water vapor. Increases in stratospheric water

vapor that are roughly twice the expected increase from

methane oxidation have now been measured by satellite

and ground-based observing systems (see Chapter 4 of

this Assessment; Rosenlof et al., 2001; SPARC, 2000).

Several recent papers have examined the consequences of

this increase for stratospheric temperatures (see Figure 3-

40d).

The study of Forster and Shine (1999) used an oversimplified representation of the water vapor change, in

both the perturbation and background stratospheric water

values (Forster and Shine, 2002) and, perhaps, in its radiative transfer scheme (Oinas et al., 2001). In general, inadequacies of broadband radiative transfer codes are readily

accounted for (see Forster et al., 2001; Forster and Shine,

2002). Since the Forster and Shine (1999) study, Smith

(2001) (see also Smith et al., 2001) used trends derived

from HALOE data over the period 1992 to 1999 to represent better the water vapor change (see also Appendix 3A

for HALOE data descriptions). Likewise, Forster and

Shine (2002) used an improved representation of the background water vapor and a simulated +1-ppmv water vapor

increase, derived from SPARC (2000) data. Shindell

(2001) also modeled the effect on increases in stratospheric water vapor, from methane oxidation, using a

coupled chemistry-climate GCM. Although large uncertainties in the water vapor trend remain, these studies

indicate a possible cooling of more than 0.5 K/decade at

70°N and 70°S, comparable with that due to ozone (Figure

3-40a). Given the present level of uncertainty in the

observational trend analysis and notwithstanding the contribution of WMGHGs and stratospheric ozone to the

cooling, one cannot rule out a significant effect due to

stratospheric water vapor. There is an indication that a

water vapor contribution is required for a quantitative

accounting of the observed polar cooling (also see the discussion in Chapter 4).

3.4.3.5 SOLAR CHANGES

Following on from work discussed in the reviews

of WMO (1999) and Ramaswamy et al. (2001), van Loon

and Labitzke (2000) correlate the solar cycle with stratospheric temperatures and find that the “response” of the

Arctic stratosphere depends on the phase of the QBO.

During solar maximum, easterly phases of the QBO coincide with a cooler Arctic stratosphere, whereas westerly

phases of the QBO give a warmer stratosphere. A model

simulation (Cubasch and Voss, 2000) was unable to simulate this response, although their model did not include a

modulation of ozone with ultraviolet radiation. Inclusion

of this feedback affects the Arctic temperature response

(Haigh, 1999; Larkin et al., 2000). Both these modeling

studies found a warmer wintertime Arctic stratosphere

during solar maximum. Recent studies with coupled

chemistry-climate models (Williams et al., 2001; Labitzke

et al., 2002) find increased sensitivity of lower stratospheric ozone to the solar cycle, compared with earlier

modeling experiments, giving greater consistency with

observations (see also EC, 2001). Generally, the latest

modeling studies indicate a possible influence of the solar

cycle on high-latitude temperatures. However, any effect

is still too uncertain to quantify and remains somewhat

speculative.

3.4.3.6 VOLCANOES

In the last 20 years the two large volcanic eruptions

of Mt. Pinatubo and El Chichón created significant

amounts of aerosol in the low-latitude stratosphere.

Although several simulations show a low-latitude stratospheric warming, the response at high latitudes is less

certain (WMO, 1999). Ramachandran et al. (2000) simulated the Mt. Pinatubo eruption in a GCM and found that

the dynamical response to the aerosol forcing led to an

POLAR OZONE

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annually averaged cooling at high latitudes of up to 2 K

for the 2 years following the eruption. Ramaswamy et al.

(2002b) find that the high-latitude simulations of Mt.

Pinatubo aerosol effects are affected substantially by the

initial conditions in the model ensemble integrations.

Timmreck et al. (1999) also found that volcanic aerosol

caused a stronger Arctic vortex. Although both model and

observations in the Mt. Pinatubo case indicate that polar

effects were not statistically significant, the suggestion

remains that volcanic eruptions may have contributed

to the observed high-latitude stratospheric cooling.

Following a volcanic eruption, enhanced ozone loss is

expected from heterogeneous chemical reactions on the

volcanic aerosol (see WMO, 1999). This ozone loss

would cool the stratosphere. Pawson et al. (1999) postulate a stepwise decrease in stratospheric temperatures following volcanic eruptions that may be connected with

changes in heterogeneous ozone chemistry (Solomon et

al., 1998).

3.4.4 Timing of the Springtime Cooling

Trend

Observational evidence shows that the maximum

springtime stratospheric cooling trend occurs at roughly

the same time as a maximum ozone loss (March in the NH

and October-November in the SH (e.g., Randel and Wu,

1999a; see also Figure 3-9). However, most modeling

studies of ozone loss and simple radiative arguments

would suggest that the maximum cooling lags the ozone

loss in the lower stratosphere by a month or more (see

Figure 3.41a,b; see also Forster et al., 1997; Ramaswamy

et al., 2001; Langematz, 2000; Rosier and Shine, 2000;

Graf et al., 1998; Langematz et al., 2002). For example,

in Antarctica a number of the models show approximately

similar cooling trends in SON and DJF for their stratospheric ozone experiments (Figure 3-40). This is especially true of the 100-hPa level, where radiative relaxation

times are longer. The NCEP reanalysis also hints at this

feature (Figure 3-41, panel c). This feature is not observed

in the satellite record. Graf et al. (1998) investigated

whether WMGHG increases could compensate for this

lag and found that they could not, at least in their model.

Ramaswamy and Schwarzkopf (2002) find that the effects

due to ozone and WMGHG over the period 1980 to 2000

in the northern polar region are swamped by the dynamical variability seen in both model and observations. The

finding is probably true of the other models (Figure 3-40).

Graf et al. (1998) further suggested that the discrepancy

could be due to an incorrect modeling of the dynamical

response and concluded that, until reasons for this are

adequately resolved, it represents a problem in the attribution, of particularly the Arctic cooling, to an anthropogenic cause.

3.4.5 Summary and Conclusions

Generally, modeling studies now demonstrate that

the springtime cooling in the Arctic lower stratosphere

over the 1980 to 2000 period is, in part, due to stratospheric ozone depletion, but the degree of attribution is

hindered by the large dynamical variability in this region.

In Antarctica, modeling studies reaffirm that ozone loss is

the major cause of the springtime cooling. They also indicate that WMGHG and stratospheric water vapor

increases are likely contributors to the annually averaged

cooling.

For combined changes in stratospheric ozone and

WMGHGs there is generally a reasonable agreement

between the various modeling studies; inter-model differences are generally smaller than the uncertainty in observations (the 1s uncertainty in the observations, if plotted

on Figure 3-39, would be substantially larger than the

plotted model error bars, which indicate the inter-model

range). Including the effects of stratospheric water vapor

in models improves the comparison with observations (the

“All” scenario of Figure 3-39). However, there is still a

large spread in model results, especially for the spring

seasons. For the annual mean temperature change, both

the agreement between models and the agreement between

the satellite and sonde-based observations are encouraging (see Figure 3-39 and Chapter 4). 

In summary the cooling of the springtime highlatitude stratosphere is likely influenced to a substantial

extent by various mechanisms (WMGHG increases,

ozone loss, stratospheric water vapor increases, volcanic

eruptions, and natural variability). Attribution of polar

temperatures is hampered by the large variability in the

polar vortices (see Figure 3-8). This variability not only

increases the uncertainty of observational trend analyses,

but it also means that either many years of equilibrium

studies or many ensemble integrations are needed for reliable statistics. Further, it implies that model studies performed under equilibrium conditions may be inadequate

for examination of polar trends, especially in the Arctic,

over a time period of two decades. At present, 20 years

of an equilibrium GCM run are barely sufficient to resolve

temperature changes in the Arctic winter. This would

imply that 20-plus-member ensembles would also be

needed for transient integrations. Comparison between

observations and models, or between different models, is

also complicated by uncertainty in the vertical profile of

ozone loss, which leads to uncertainty in the temperature

response. In models that simulate the chemistry, this

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uncertainty could be larger than in those models that prescribe the ozone loss from observations. These coupledchemistry models are discussed in the next section.

3.5 CHEMICAL-CLIMATE MODELING OF

THE PAST AND FUTURE POLAR

STRATOSPHERE

3.5.1 Introduction

Section 3.3.2 shows that, while there has been

halogen-induced chemical ozone loss in many recent

Arctic winters, the extent of this chemical loss in a given

year is determined by the meteorological conditions, with

severe ozone loss in the colder winters and essentially no

loss in the warmer winters. Thus, low temperatures are

required to initiate wintertime chemical ozone loss in the

Arctic, and the time series of Arctic ozone shown in Figure

3-2 reflects the changes in meteorology much more than

it does the changes in halogen loading. This understanding is consistent with the conclusion of Section 3.4.3

(see also Rosier and Shine, 2000; Langematz, 2000) that

the observed decrease in Arctic temperature over the last

two decades cannot be explained just as a radiative

response to the observed ozone decrease (although there

is a contribution), and must have occurred for some other

reason. In this respect the Arctic is quite different from

the Antarctic. In the Antarctic, wintertime temperatures

are low enough to initiate severe chemical ozone loss

every year, and meteorological variability provides only a

small modulation of this loss. The time series of Antarctic

ozone shown in Figure 3-2 therefore primarily reflects

changes in halogen loading, and the observed late-spring

cooling and increased persistence of the Antarctic vortex

are understood to be a radiative response to the observed

ozone decrease (Section 3.4.3). That this Antarctic

cooling is less than the observed Arctic cooling, despite

the fact that Antarctic ozone losses are much greater than

those in the Arctic, provides further confirmation that the

Arctic cooling is a driver of the Arctic ozone decreases

(in the presence of anthropogenic halogens), rather than

being a response to them.

The difference between the past behavior of Arctic

and Antarctic ozone has implications for the future evolution of polar ozone. In the Antarctic, meteorological conditions are so far past the threshold for severe chemical

ozone loss that neither natural meteorological variability

nor WMGHG-induced climate change are expected to

change this situation over the coming decades. Thus,

future Antarctic ozone levels are expected primarily to

reflect the anticipated decrease in stratospheric halogen

loading, as they have reflected the observed increase in

halogen levels in the past. (There is the possibility of more

sensitivity for aspects of Antarctic ozone that are closer to

threshold conditions, such as ozone losses at the edge of

the vortex or in early summer; these may be of more

importance for UV impacts than minimum column ozone

amounts.) In the Arctic, on the other hand, our expectation for future ozone changes depends on our attribution

of the past changes, and in particular on the cause of the

observed Arctic cooling over the past two decades. If this

cooling represents a forced response to WMGHG

changes, then we would expect it to continue in the future

and Arctic ozone levels would be expected to decrease

further in spite of the decreasing halogen loading. If, on

the other hand, the observed cooling represents natural

variability, then anything is possible in the future; chemical ozone loss in the Arctic could be essentially zero for

the next decade, or it could be severe.

The implication of the above considerations is that

we cannot expect future Arctic ozone abundances to

follow stratospheric halogen loading, apart from the

obvious fact that halogen-induced chemical ozone loss

can be expected to become negligible when halogen levels

become negligible. Instead, the evolution of Arctic ozone

abundances (which reflect both chemistry and transport)

will be determined primarily by the evolution of meteorological conditions. In the Arctic, those meteorological

conditions are largely determined by dynamical processes

(Section 3.2.1), and the effects of WMGHG changes on

the lower stratospheric Arctic vortex over the past and

through the next few decades are expected to mainly arise

from dynamical feedbacks, rather than from direct radiative forcing (Section 3.4). Therefore, in order both to

attribute past behavior and predict future behavior of

Arctic ozone, it is necessary to use models that are capable

of representing those dynamical feedbacks together with

a prognostic representation of ozone, namely, coupled

chemistry-climate models.

It is important to appreciate what such models can

and cannot do. In this respect there is a notable contrast

with the CTMs used in Section 3.3. CTMs specify the

meteorological conditions (winds and temperatures), and

predict the resulting chemical ozone loss without

addressing the issue of why those conditions occurred.

This prediction is essentially deterministic, because the

chemical equations are essentially deterministic; we

expect to capture the exact chemical ozone loss on a particular day, within the uncertainties of the calculation.

Coupled chemistry-climate models specify the radiative

and chemical forcing (WMGHGs, halogens, surface conditions) and predict the resulting state of the chemistryclimate system. However, since the climate system is

chaotic and exhibits unforced variability, this prediction

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is only stochastic, not deterministic; we do not expect to

capture the exact state of the chemistry-climate system on

a particular day, or even in a particular year. Therefore,

the comparison between such models and observations

requires careful thought. On the other hand, chemistryclimate models allow one in principle to draw unambiguous conclusions about the attribution of past changes,

and they are the only tool available for prediction of future

changes that includes all physical feedbacks.

There are two distinct strategies for assessing the

response of a chemistry-climate model to chemical or

radiative forcing. The traditional approach is to perform

simulations under fixed conditions (so-called “timeslice”

experiments) and compare the differences in equilibrium

climate resulting from the different forcings. The advantage of this approach is that the simulations can, in principle, be run for a sufficiently long time to obtain good

statistics on the differences, which is useful for understanding mechanisms and feedbacks. In this way the

unforced variability inherent in the climate system is dealt

with explicitly, albeit in an idealized manner. The second

approach is to perform simulations under changing conditions (so-called “transient” experiments) over a particular period of time. The advantage of this approach is that

it permits a more direct comparison with the real atmosphere, where the forcings are evolving continuously in

time. Because of unforced variability, two realizations of

the same transient experiment will not evolve in exactly

the same way, and an ensemble of experiments is generally required in order to determine a statistically reliable

signal.

Yet no matter how good a chemistry-climate model

is and how well its climate-change experiments are characterized, the real atmosphere exhibits only a single realization. The best one can expect, even for a perfect model,

is that the observations fall within the range of modelpredicted behaviors, according to appropriate statistical

criteria. Whether this permits a meaningful prediction

depends on the relative magnitudes and time scales of the

forced signal and the natural noise. As discussed above,

we expect the evolution of Antarctic ozone over decadal

time scales (past and future) to be a fairly predictable

signal, but it is not at all clear whether this is the case in

the Arctic. It is known that there is strong interannual

variability in the Arctic, extending to multidecadal time

scales; even timeslice experiments require several decades

of integration to obtain reliable statistics for basic diagnostic quantities (Scaife et al., 2000b). Thus, it is not a

priori obvious that even a perfect chemistry-climate model

should be able to reproduce the long-term decreases of

Arctic ozone observed over the past 20 years, for example.

Chemistry-climate models for stratospheric ozone

are probably the most advanced of all chemistry-climate

models, but they are still being developed. Early timeslice experiments were performed by Austin et al. (1992)

and Pitari et al. (1992) to examine the question of how

Arctic ozone might evolve under the conditions of doubled CO2 expected toward the end of the 21st century, if

large halogen abundances were to persist until that time.

The results were contradictory, with the former study predicting severe Arctic ozone loss and the latter study predicting only a slight reduction; this at least demonstrated

the sensitivity of the response. In the previous Assessment

(WMO, 1999, Chapter 12), many of the conclusions were

based on the transient experiment of Shindell et al.

(1998a), which examined the combined effects of anticipated increasing WMGHGs and decreasing halogens, and

predicted increasingly severe Arctic ozone loss over the

next couple of decades resulting from Arctic cooling. The

stated mechanism for this severe cooling was the dynamical feedback from WMGHG changes, and this represents

the key question for the Arctic: namely, are the dynamical

feedbacks resulting from WMGHG-induced temperature

changes expected to enhance or to reduce the direct radiative stratospheric cooling?

Since the previous Assessment, a number of coupled chemistry-climate models have been developed and

run, so there is now considerably more information available than at that time (see Austin et al., 2002). Although

these models have improved representations of many

processes, they still have many deficiencies, and computational constraints also place severe limitations on the

experimental configurations; transient experiments are

generally performed singly, not as part of an ensemble,

and timeslice experiments are often too short to adequately characterize their variability. In this report, the

strategy is therefore adopted of combining the results of

the different models, whether the experiments are timeslice or transient, the working assumption being that the

differences between the models cover the range of model

errors. In this way a current best estimate is obtained of

the attribution of past ozone changes in the polar stratosphere and the prediction of future changes over the next

half-century.

The models used in the comparisons are indicated

in Table 3-3, in order of decreasing horizontal resolution

(see Austin et al., 2002). Università degli Studi

dell’Aquila (ULAQ) is the only model with a substantial

aerosol package and it has reasonably detailed chemistry,

albeit diurnally averaged. This model has been run in

timeslice mode. Of the other models run in this mode, the

Canadian Middle Atmosphere Model (CMAM) and the

Middle Atmosphere European Centre Hamburg Model

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(ECHAM) with Chemistry (MAECHAM/CHEM) have

reasonably detailed chemistry and a high upper boundary

(0.01 hPa and above), while the University of Illinois at

Urbana-Champaign (UIUC) model and the ECHAM

model with chemistry run at DLR (E39/C) have a much

lower upper boundary (1 hPa and below). The Unified

Model with Eulerian Transport and Chemistry (UMETRAC), Center for Climate System Research (University

of Tokyo)/National Institute for Environmental Studies

(Japan) (CCSR/NIES), and GISS have been run in transient mode; UMETRAC and CCSR/NIES have reasonably detailed chemistry, whereas the GISS model has

parameterized chemistry.

3.5.2 The Uncertainties in ChemistryClimate Models

3.5.2.1 TEMPERATURE BIASES

As noted in Section 3.4.2.2, many climate models

without chemistry but with a fully resolved stratosphere

have a cold bias of the order of 5 to 10 K over Antarctica

in the lower stratosphere. This suggests that their residual

circulations are too weak (Pawson et al., 2000), i.e., there

is too little downwelling in high latitudes and too little

upwelling in low latitudes. This cold bias would significantly impact model heterogeneous chemistry and

enhance ozone destruction. The “cold-pole problem”

extends to higher stratospheric levels, causing a polar

night jet that is too strong and too vertically oriented,

whereas the observed polar night jet slopes with height

toward the equator in the upper stratosphere. The weaker

jet and vertical slope allow waves to propagate into higher

latitudes and maintain higher polar temperatures. A potentially important component of climate change is whether

these waves will be stronger in the future, since this will

likely affect the evolution of ozone (see Section 3.5.2.4).

A practical solution applied in some models with a cold

bias is to increase the temperatures in the heterogeneous

chemistry routines (e.g., Austin et al., 2000) by a fixed

amount (e.g., 5 K). The strong polar night jet is also associated with a polar vortex that breaks down later in the

spring, particularly in the SH. In a chemistry-climate

model this can lead to ozone depletion that continues for

longer than observed.

Gravity waves generated by processes other than

orographic forcing (e.g., convection) are thought to play

an important role in the momentum balance of the stratosphere. Non-orographic gravity wave drag (gwd) schemes

have now been developed for climate models (e.g.,

Medvedev and Klaassen, 1995; Hines, 1997a, b; Warner

and McIntyre, 1999), and their use has resulted in a significant reduction in the cold-pole problem relative to simulations that rely on Rayleigh friction to decelerate the

POLAR OZONE

Table 3-3. Models used in the comparisons. The resolution is given in either degrees latitude ¥ degrees

longitude (grid point models), or as T21, T30, and T32, which are the resolutions in spectral models corresponding to triangular truncation of the spectral domain with 21, 30, and 32 wave numbers, respectively.

IS92a refers to scenario IS92a of IPCC (1992), and WMO refers to the halogen scenario indicated in Chapter

11 of WMO (1999).

Model Horizontal No. Levels/ GHG/ Reference

Resolution Upper Halogen

Boundary Scenarios

UMETRAC 2.5° ¥ 3.75° 64/0.01 hPa IS92a/WMO Austin (2002) (Rayleigh friction version)

CMAM T32 65/0.0006 hPa Observations, de Grandpré et al. (2000)

IS92a/WMO

MAECHAM/ T30 39/0.01 hPa IS92a/WMO See details in Austin et al. (2002)

CHEM

E39/C T30 39/10 hPa IS92a/WMO Schnadt et al. (2002)

UIUC 4° ¥ 5° 25/1 hPa Observations Rozanov et al. (2001)

CCSR/NIES T21 30/0.06 hPa IS92a/WMO Takigawa et al. (1999);

Nagashima et al. (2002)

GISS 8° ¥ 10° 23/0.002 hPa IS92a/WMO Shindell et al. (1998b)

ULAQ 10° ¥ 20° 18/1 hPa IS92a/WMO Pitari et al. (2002a)

3.68
polar night jet (e.g., Manzini and McFarlane, 1998). Two

of these schemes have also been shown to produce a QBO

when run in a climate model (Scaife et al., 2000a). The

latest versions of several coupled chemistry-climate

models now employ such schemes: CMAM uses the

Medvedev-Klaassen scheme (Medvedev et al., 1998) or

the Hines scheme (McLandress, 1998); UMETRAC uses

the Warner and McIntyre scheme; and MAECHAM/-

CHEM uses the Hines scheme (see Austin et al., 2002).

The GISS GCM has used a non-orographic gravity wave

drag scheme for many years (Rind et al., 1988a, b), which

is able to reproduce high-latitude temperatures reasonably well (Shindell et al., 1998b) but does not simulate a

QBO in the tropics.

Figure 3-42 shows model temperature biases as a

function of height for 80°N and 80°S, which are representative of the polar regions, for the winter and spring

seasons. To determine the biases, a 10-year temperature

climatology determined from UKMO data assimilation

fields (Swinbank and O’Neill, 1994) was subtracted from

the mean model temperature profiles applicable to the

1990s. The UKMO temperatures are considered to be

about 2 K too high at low temperatures (e.g., Pullen and

Jones, 1997; Manney et al., 2002; Pommereau et al., 2002;

Knudsen et al., 2002), but this bias is smaller than typical

model biases. The upper stratospheric cold-pole problem

is particularly noticeable in the south in the UMETRAC

(with Rayleigh friction), CCSR/NIES (which also uses

Rayleigh friction), and MAECHAM/CHEM results. In

the results of the UIUC and ULAQ models, a warm bias

is present. For the ULAQ model this is likely to be due to

the inclusion of vertical diffusion (in addition to Rayleigh

friction). As seen in the UMETRAC results, the winter

and spring polar temperature bias can be dramatically

reduced by the use of non-orographic gwd. Both UMETRAC and CMAM have very similar biases, within a

few K of each other in the seasons investigated. The

MAECHAM/CHEM model, which uses the Hines nonorographic gwd scheme, is only a slight improvement

on the Rayleigh friction results of UMETRAC and

CCSR/NIES. The E39/C model is very similar to the

MAECHAM/CHEM model, and gives similar results

below 30 hPa, but does not have a non-orographic gwd

scheme. At 80°N, temperature biases are somewhat

smaller than at 80°S and are sometimes positive. The

northern lower stratospheric temperature biases would

generally lead to insufficient heterogeneous ozone depletion in early winter but excessive ozone depletion in the

more important spring period. In the SH, spring cold

biases could lead to more extensive PSCs than observed

and delayed recovery in Antarctic ozone.

Some indication of the source of the model temperature biases in the lower stratosphere is given by Figure

3-43, which shows the zonal mean heat flux [v¢T¢] at 100

hPa averaged over the domain 40° to 80°N for January

and February plotted against temperature averaged at 50

hPa over the domain 60° to 90°N for February and March.

As argued by Newman et al. (2001), the heat flux at 100

hPa is indicative of the wave forcing from the troposphere,

and this is highly correlated with lower stratospheric temperature slightly later in the year. Newman et al. (2001)

chose a 1-15 March temperature average, but here we

choose a longer period for the temperature average to

smooth model and atmospheric transients. The model

results (Figures 3-43 and 3-44) are in general agreement

with observations. Table 3-4 shows results of the linear

regression between the two variables (see table caption

for explanation of the terms T0 and b). The results indicate that horizontal resolution may have affected the

model results; in general the model regression lines were

less steep (smaller b in Table 3-4) as the model resolution

decreased, particularly in the NH. This could be because

low-resolution models capture the low-amplitude-wave,

small-heat-flux case, but have difficulty capturing the

large-heat-flux case with its significant potential

enstrophy cascade to larger wave numbers. The performance of the models might also depend on the dissipation

that the models have at short spatial scales, although this

is more difficult to compare. The values of b are generally much smaller in the SH, except for the CCSR/NIES

and CMAM models. The implication of these results is

that models need higher resolution and non-orographic

gwd schemes to improve the relationship between heat

flux and temperature in order to reduce their polar temperature biases.

3.5.2.2 THE SIMULATION OF POLAR STRATOSPHERIC

CLOUDS

PSCs have a significant impact on stratospheric

ozone depletion in polar regions, and recent developments

in their understanding are discussed in detail in Carslaw

et al. (2001) and in Section 3.2.2. Coupled chemistryclimate models have a variety of PSC schemes with and

without sedimentation, but such models have in some

cases large climatological temperature biases in the polar

regions, as indicated in Section 3.5.2.1. If the models are

to be effective, the temperature field must give realistic

distributions near the PSC temperature thresholds.

The areal coverage of PSCs provides a model comparison diagnostic. We use the temperature at 50 hPa as

an indicator of likely PSC amounts, and ignore the effect

of HNO3 and sulfate concentrations on the determination

POLAR OZONE

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of PSC surface areas (following Pawson and Naujokat,

1997; and Pawson et al., 1999). Figure 3-45 shows for

the models and observations the time integral throughout

the winter of the PSC area at 50 hPa (At). At is the sum

throughout the winter of the PSC area (expressed as %

fraction of the hemisphere) at 50 hPa. For the ice amount,

At varies dramatically in the Arctic, between zero and

70% of the hemisphere times days. The models have large

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Dec/Jan/Feb 80°N

-30 -20 -10 0 10

Model Temperature bias (K) 

1000

300

100

30

10

3 Pressure hPa

Mar/Apr/May 80°N

0 10

Model Temperature bias (K) 

1000

300

100

30

10

3 Pressure hPa

Jun/Jul/Aug 80°S

0 10

Model Temperature bias (K) 

1000

300

100

30

10

3 Pressure hPa

Sep/Oct/Nov 80°S

0 10

Model Temperature bias (K) 

1000

300

100

30

10

3 Pressure hPa

_ _ _ UMETRAC Rayleigh friction ____ UMETRAC Non-orographic GWD ____ CMAM ____ MAECHAM/CHEM ____ E39/C____ CCSR/NIES _ _ _ UIUC ____ ULAQ _ _ _ GISS

-30 -20 -10

-30 -20 -10 -30 -20 -10

Figure 3-42. Model temperature biases at 80°N and 80°S for the winter and spring seasons, as a function of

pressure. To determine the bias, a climatology determined from 10 years of UKMO data assimilation temperatures was subtracted from the model results.
interannual variability. Arctic NAT also covers a large

range, both for different models and in the interannual

variability for each simulation. In accordance with their

temperature biases, several models have larger areas of

NAT than are typically derived from observations. The

ULAQ PSCs are in good agreement with observations in

both hemispheres, despite a slight positive temperature

bias, whereas UMETRAC and CMAM have lower PSC

areas in the Arctic than are derived from observations. In

the Antarctic each model has much lower fractional interannual variation for both NAT and ice, but again the results

for separate models cover an exceedingly large range for

the ice amount. Clearly, the differences between different

models will have a profound impact on the amount of

chemical ozone depletion calculated. This will be discussed further in Section 3.5.3.

POLAR OZONE

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0 5 10 15 20 25

Heat flux at 100 hPa, Jan and Feb mean

200

205

210

215

220

225 Mean Temperature Feb and March (a) (b)

(c)

0 5 10 15 20 25

Heat flux at 100 hPa, Jan and Feb mean

200

205

210

215

220

225 Mean Temperature Feb and March

 (a) Observations

(a)

 (b) UMETRAC Non-orographic gwd

 (c) CCSR/NIES

 (f) ULAQ

 (g) CMAM 2000 run

 (d) MA-ECHAM CHEM 1990 run

(d)

 (e) E39/C

(e)

0 5 10 15 20 25

Heat flux at 100 hPa, Jan and Feb mean

200

205

210

215

220

225 Mean Temperature Feb and March (a)

(f)

(g)

Figure 3-43. Scatter diagrams of the longitudinally averaged (zonal-mean) heat flux (K m/s, averaged for 40°

to 80°N, at 100 hPa for January and February) versus temperature (K, averaged for 60° to 90°N, at 50 hPa for

February and March) for participating models. The solid lines are linear regression lines between the two variables. The observations are taken from NCEP data.
3.5.2.3 THE POSITION OF THE UPPER BOUNDARY OF

THE GCM

There is strong evidence from a number of modeling studies (Garcia and Boville, 1994; Shepherd et al.,

1996; Lawrence, 1997; Austin et al., 1997; Rind et al.,

1998; Beagley et al., 2000) that the position of the model

upper boundary can play a significant role in influencing

transport and stratospheric dynamics due to the “downward control principle” (Haynes et al., 1991). The sensitivity of the dynamical fields to the position of the upper

boundary may be larger when non-orographic gwd

schemes are used than when Rayleigh friction is used.

However, if all the non-orographic gwd that is produced

above the model boundary is placed instead in the top

model layer, assuming that the upward-propagating waves

are not simply absorbed in the top layer, this sensitivity

reduces (Lawrence, 1997). Model simulations with an

upper boundary as low as 10 hPa have been completed

(e.g., Schnadt et al., 2002; Hein et al., 2001; Dameris et

al., 1998). Schnadt et al. (2002) show the meridional

circulation of the Deutschen Zentrum für Luft- und

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0 2 4 6 8 10 12

Heat flux at 100 hPa, Jul and Aug mean

185

190

195

200

205 Mean Temperature Aug and Sept (a)

(b)

(c)

0 2 4 6 8 10 12

Heat flux at 100 hPa, Jul and Aug mean

185

190

195

200

205 Mean Temperature Aug and Sept (a)

(d)

(a) Observations

(b) UMETRAC Non-orographic gwd

(c) CCSR/NIES

(g) CMAM 2000 run

(d) MA-ECHAM CHEM 1990 run

(e) E39/C

(e)

0 2 4 6 8 10 12

Heat flux at 100 hPa, Jul and Aug mean

185

190

195

200

205 Mean Temperature Aug and Sept (a)

(g)

Figure 3-44. As in Figure 3-43, but for heat flux at 40° to 80°S, at 100 hPa for July and August, and temperature at 60° to 90°S, at 50 hPa for August and September.
Raumfahrt (DLR) model. This gives the expected upward

motion from the summer hemisphere and downward

motion over the winter hemisphere, although modeled

meridional circulations are known to extend into the

mesosphere (e.g., Butchart and Austin, 1998). Schnadt et

al. (2002) argue that it is less important to have a high

upper boundary, but more important to have high resolution in the vicinity of the tropopause. At present the evidence appears ambiguous; for example, in the total ozone

presented by Hein et al. (2001), insufficient ozone is

transported to the North Pole, but there is excessive subtropical ozone transport. This could be related to the

cold-pole problem, rather than the position of the upper

boundary. While the transport effect on ozone is direct,

other considerations are the transport of long-lived tracers

such as NOy and water vapor that have a photochemical

impact on ozone. Consequently, it is generally recognized

that the upper boundary should be placed at least as high

as 1 hPa (e.g., Rozanov et al., 2001; Pitari et al., 2002b);

many models are now placing their boundary at about

0.01 hPa (e.g., Shindell et al., 1998b; Austin et al., 2001;

Nagashima et al., 2002). In comparison, CMAM (de

Grandpré et al., 2000) has an upper boundary somewhat

higher (~0.0006 hPa) to allow a more complete representation of gwd to reduce the cold-pole problem (Section

3.5.2.1) and to simulate upper atmosphere phenomena.

3.5.2.4 PREDICTIONS OF PLANETARY WAVES

The direct radiative cooling by greenhouse gases

at high latitudes in the lower stratosphere causes an

increase in the strength of the polar vortex. Planetary

wave changes may be a feedback that strengthens or

weakens this effect. One proposed planetary wave feedback mechanism (Shindell et al., 2001) works as follows:

tropical and subtropical sea surface temperatures increase,

leading to a warmer tropical and subtropical upper troposphere via moist convective processes. This results in an

increased latitudinal temperature gradient at around 100

to 200 hPa, leading to enhanced lower stratospheric westerly winds that refract upward, propagating tropospheric

planetary waves equatorward. This results in a strengthened polar vortex.

The E39/C results (Schnadt et al., 2002) also show

a small decrease in planetary wave activity for the period

1960 to 1990. However, results for future simulations

differ depending on the model. Without chemical feedback, the Unified Model (UM) (Gillett et al., 2002) predicts a future increase in overall generation of planetary

waves, which leads to a greater wave flux to the Arctic

stratosphere, and is even able to overcome the radiatively

induced increase in the westerly zonal wind so that the

overall trend is to weaken westerly flow. This also occurs

in the DLR model with chemical feedback (Schnadt et al.,

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Table 3-4. Statistical analysis of the linear regression between the area-averaged temperature (K) at

50 hPa poleward of 60°N for February and March in the NH and 60°S in August and September in the

SH, and the heat flux (K m/s) at 100 hPa between 40° and 80°N for January and February in the NH and

40° and 80°S for July and August in the SH. R is the correlation coefficient between the variables, T0 is the

intercept of the line at zero heat flux, and b is the gradient of the line. For ULAQ results in the Southern

Hemisphere, the heat flux and temperature are not correlated.

Model/Observations Northern Hemisphere Southern Hemisphere

R T0 b R T0 b

NCEP (observations) 0.77 195.1 1.49 0.78 189.4 0.89

UMETRAC non-orographic gwd 0.74 196.9 1.21 0.66 188.5 0.98

Rayleigh friction 0.67 196.2 1.21 0.51 187.7 0.67

CMAM 2000 0.52 204.5 0.75 0.44 187.2 0.68

MAECHAM/CHEM 1990 0.79 196.3 1.10 0.70 190.0 0.50

E39/C 0.62 198.3 0.93 0.86 186.3 0.56

CCSR/NIES 0.74 199.4 0.80 0.74 186.6 1.17

ULAQ 0.58 203.0 0.48 —— —— ——
2002), whereas in the UM with chemical feedback (UMETRAC) the trend in the heat flux during the period 1975

to 2020 is downward but is not statistically significant.

The Center for Climate System Research/National

Institute for Environmental Studies (CCSR/NIES) model,

of lower resolution than UMETRAC, has systematically

lower heat fluxes but does show a downward trend during

the period 1986 to 2050 that is marginally statistically significant. In general, the strengthening of the polar vortex

appears to be critically dependent upon the relative importance of changes in wave generation versus wave propagation. These changes are likely to be highly model and

POLAR OZONE

3.74

Arctic NAT

1960 1980 2000 2020

0

200

400

600 Seasonal PSC % days

Arctic ICE

1960 1980 2000 2020

0

50

100

150 Seasonal PSC % days

Antarctic NAT

1960 1980 2000 2020

0

500

1000

1500

2000 Seasonal PSC % days

Antarctic ICE

1960 1980 2000 2020

0

200

400

600

800

1000 Seasonal PSC % days

____ _ _ _ UMETRAC Non-orographic GWD

E39/C

CMAM

Observations

MAECHAM/CHEM

ULAQ

Figure 3-45. Time integral of the percentage area of the hemisphere at 50 hPa covered by temperatures

below 195 K (approximate NAT temperature) and below 188 K (approximate ICE temperature) during winter

and spring. The error bars for the timeslice runs indicate 95% confidence intervals. For clarity, ULAQ data are

plotted 2 years early, and MAECHEM CHEM data for 1990 and 2030 are plotted 2 years late.
resolution dependent, resulting from the particular wave

forcing and drag schemes employed in each climate

model. To some degree this sensitivity of the changes in

planetary waves to the model simulation reflects similar

sensitivities in the atmosphere: for example, in the observations of the last five northern winters, four have been

warm with a weaker polar vortex (see Section 3.1.2 and

Figure 3-8). As a result, the strengthening of the vortex

over the last 20 years noted by some authors (Tanaka et

al., 1996; Zurek et al., 1996; Waugh et al., 1999; Hood et

al., 1999) will have been modified by recent events.

3.5.2.5 STRATOSPHERIC WATER VAPOR CHANGES

Observations of atmospheric water vapor concentrations have revealed significant increases over the period

1964 to 2000 (Oltmans and Hofmann, 1995; Oltmans et

al., 2000; Rosenlof et al., 2001). As discussed further in

Chapter 4, these increases are uncertain in magnitude and

their causes have not been established.

Increased water vapor directly affects ozone chemistry. It also alters local temperatures by radiative cooling,

slowing down the reaction rates of ozone depletion chemistry, which indirectly leads to more ozone. Evans et al.

(1998), Dvortsov and Solomon (2001), and Shindell

(2001) have studied the effects on homogeneous chemistry. The models all show that increases in water vapor

reduce ozone in the upper and lower stratosphere, and

increase ozone in the middle stratosphere. The model

results differ most in the lower stratosphere where the

largest impact on total ozone column occurs. In the

model of Evans et al. (1998), lower stratospheric ozone is

reduced only in the tropics when water vapor increases,

whereas in the other models, ozone reductions extend to

midlatitudes or to the poles. Thus, the models of Dvortsov

and Solomon (2001) and Shindell (2001) show a slower

ozone recovery by about 10 to 20 years, and a 1-2% reduction during the next 50 years due to water vapor increase,

assuming that the water vapor increase continues at the

current rate. The Evans et al. (1998) model disagrees with

these results, presumably due to differences in that

model’s temperature response to increasing water, which

seems to dominate over the chemical impacts.

Water vapor increases also affect heterogeneous

chemistry, enhancing the formation of PSCs (see Section

3.2.2). Kirk-Davidoff et al. (1999) calculated a significant enhancement to Arctic ozone loss in a more humid

atmosphere. Much of this effect was based on a very large

estimate of 6 to 9 K per ppmv radiative cooling induced

by increased water vapor. This value has been superseded

by newer results showing that the value is almost certainly

much smaller, about 1.5 to 2.5 K cooling per ppmv of

water (Section 3.4.3.4). This would in turn imply a

reduced role for water vapor in enhancing PSC formation.

Tabazadeh et al. (2000) showed that the enhancement of

PSC formation due to the addition of 1 ppmv of water

vapor is approximately the same as the PSC enhancement

due to cooling of about 1 K. This suggests that the radiative impact of water vapor is larger than its effects on

chemistry or microphysics but that all these processes

should be considered in numerical models. Given the

potential for denitrification in the Arctic, and the large

ozone losses that could result from a slight cooling there

(Tabazadeh et al., 2000), it is important both to understand

trends in stratospheric water vapor and to resolve model

differences in the radiative impact of those trends.

Model simulations of past water vapor trends do not

agree well with observations. In UMETRAC (Austin,

2002), water vapor increases by only about 1% per decade

in the stratosphere, despite the inclusion of a methane

oxidation scheme. In UMETRAC, the tropical tropopause

temperature decreases slightly, counteracting the methane

impact. In E39/C (Schnadt et al., 2002), water vapor

increases in the lower stratosphere are significantly larger

(about 3% per decade) but are still about a factor of 2 to 3

lower than observed. Similar results are also obtained in

the GISS model. In general the modeled water vapor trend

tends to be driven by methane oxidation and trends in tropical tropopause temperatures, suggesting the need to investigate the microphysics of dehydration and how this is represented in models. (See also Section 4.3.3 of EC (2001).)

3.5.3 Model Assessments

In the Arctic the processes leading to stratospheric

ozone depletion may undergo too much natural variability

to provide a definite answer to how ozone will actually

evolve. Each model may be considered as supplying a

single simulation (or range of simulations in the case of

the timeslice experiments) of a larger ensemble. Although

the mean of the ensemble can be readily computed, the

atmosphere may in practice evolve in a manner anywhere

within, or even outside, the envelope of the model simulations. In the Antarctic, the dominant processes are less

dependent on interannual variability and, hence, the ozone

evolution is in principle more predictable.

One of the emphases here has been on spring ozone

recovery. In view of the range of results obtained, it is important to define this term carefully. It is here used in two

senses: (1) the start of ozone recovery, defined as the date of

the minimum spring column ozone as a function of year in

the decadally averaged results, and (2) full ozone recovery,

defined as the date of the return of the decadally averaged

spring column ozone to the value obtained in 1980.

POLAR OZONE

3.75
POLAR OZONE

3.76

(a) Transient runs

1960 1970 1980 1990 2000 2010 2020 2030 2040 2050 2060

100

150

200

250

300

350 Min. Total Ozone (DU) TOMS

UMETRAC

CCSR/NIES

GISS - 1998 Assessment

(b) Timeslice runs

1960 1970 1980 1990 2000 2010 2020 2030 2040 2050 2060

200

250

300

350

400

450 Min. Total Ozone (DU) TOMS

CMAM

MAECHAM/CHEM

E39/C

Figure 3-46. Minimum Arctic (March/April) total ozone for the main experiments of this Assessment. (a)

Transient runs in comparison with TOMS data. The solid lines show the results of a Gaussian smoothing

applied to the individual year’s results. The error bars denote twice the standard deviation of the individual

years from the smoothed curve. Results from the previous Assessment (the GISS model) are included for

comparison. (b) Timeslice runs in comparison with TOMS data. The colored circles and the error bars denote

the mean and twice the standard deviation of the individual years within each model sample (10 years for

CMAM, 20 years for MAECHAM/CHEM and E39/C). Dotted lines are drawn between the end points of the

error bars to assist in estimating trends by eye. For MAECHAM/CHEM only: (1) the values have been plotted

2 years late for clarity, and (2) a standard tropospheric column of 100 DU has been added to the computed

stratospheric columns.
3.5.3.1 THE 1960-2000 TIME FRAME: OZONE

DEPLETION

As is well established from observations (Section

3.1.1), polar ozone has been decreasing over the last few

decades. Figure 3-46 shows the minimum daily ozone

throughout the range 60° to 90°N for several of the models

of Table 3-3 together with TOMS data. The GISS results

are included to provide a comparison with the state of

knowledge at the time of the previous Assessment (WMO,

1999). For clarity, the models are separated into two

panels according to their mode of simulation (transient or

timeslice). For the transient models, a Gaussian low-pass

filter is fitted through the results for the individual years,

and the error bars denote two standard deviations from

the Gaussian filtered curve. For the timeslice experiments, the mean and two standard deviations are plotted,

and the dotted lines give an indication of the temporal

behavior of the model results. Each model has a large

interannual variability, similar to that of the observations,

and, hence, detecting a signal is difficult. All the models

indicate a slight high bias relative to observations. The

trends in the minimum (Table 3-5) are consistent with the

observations, although only E39/C, MAECHAM/CHEM,

and the observations have statistically significant trends.

Averaging the five model results obtained since the previous Assessment gives an Arctic trend that is statistically

significant and agrees with observations.

In the Antarctic (Figure 3-47, upper panel), the

model runs agree reasonably well with observations for

the past and show the steady development of the ozone

hole during the period. However, the calculated trends

(Table 3-5) depend on the period chosen. CMAM indicates a small trend because the ozone hole was already

partially formed in 1987 and the CMAM values for 2000

are higher than observed. The mean trend of the latest

five model simulations (Table 3-5) is slightly lower than

observed, but if the latest reaction rates for the ClO dimer

production (Bloss et al., 2001) had been used, the agreement would have been further improved. While the interannual variability of most of the models is similar to that

observed, both CMAM and UMETRAC have a large

interannual variability. In the case of UMETRAC, this

may to some extent be a product of the non-orographic

gwd scheme.

The maximum size of the Antarctic ozone hole

during each spring, as given by the area within the 220-

DU total ozone contour, is shown in Figure 3-47 (lower

panel). The results for GISS, E39/C, and CCSR/NIES are

in good agreement with observations, but may indicate a

slight underprediction. A much smaller ozone hole is simulated by UMETRAC, but recent model runs with a

revised NOy distribution consistent with observations

show an ozone hole about 50% larger for 1995, which is

in close agreement with observations. The size of the

ozone hole area for CMAM reflects the bias and large

interannual variability noted in Figure 3-47 (upper panel).

Errors in the modeling of the size of the ozone hole can

have important implications. First, comparisons between

models and observations for ozone amounts near 60°S

will give poor agreement if the ozone hole area is too

small, even though the underlying physics of the model

may be correct. Second, a model with a smaller ozone

hole may evolve differently from that of the atmosphere

due to transport and chemistry effects relating to radiative

effects.

3.5.3.2 THE 2000-2020 TIME FRAME: START OF

OZONE RECOVERY

The first signs of ozone recovery are expected

within the next two decades (Shindell et al., 1998a; Austin

et al., 2000; Schnadt et al., 2002; Rosenfield et al., 2002;

Nagashima et al., 2002). Two-dimensional (2-D) model

simulations (e.g., Rosenfield et al., 2002) indicate a slight

delay in Arctic and Antarctic spring ozone recovery following the maximum values in halogen loading. The

POLAR OZONE

Table 3-5. Past trends (1979-2000) in minimum

ozone (DU/decade) with 2s error bars for participating models and TOMS. The MAECHAM/CHEM

results cover the period 1960-2000 (Arctic) and

1960-1990 (Antarctic); the E39/C (ECHAM model

with chemistry run at DLR) results cover the period

1960-1990 (Arctic) and 1980-1990 (Antarctic); the

CCSR/NIES model results cover the period 1986-

2000; and the CMAM results cover the period 1987-

2000.

Model/Observations Northern Southern

Trend Trend

UMETRAC -6 ± 22 -80 ± 31

CMAM -13 ± 27 -23 ± 28

MAECHAM/CHEM -16 ± 10 -45 ± 3

E39/C -16 ± 14 -64 ± 7

CCSR/NIES -33 ± 38 -41 ± 21

Mean (this Assessment) -17 ± 12 -51 ± 11

GISS (1998 Assessment) -21 ± 34 -34 ± 12

TOMS -21 ± 16 -59 ± 12

3.77
POLAR OZONE

3.78

Minimum Antarctic Ozone September-November

1960 1970 1980 1990 2000 2010 2020 2030 2040 2050 2060

Year

50

100

150

200

250

300 Min. Total Ozone (DU)

Maximum Modeled Antarctic Ozone Hole Area

1960 1980 2000 2020 2040 2060

Year

0

5

10

15

20

25

30

35 Area inside 220 DU (106 km2

)

TOMS

CMAM

UMETRAC

MAECHAM/CHEM

E39/C

CCSR/NIES

UIUC

ULAQ

GISS 98 Assessment

Figure 3-47. Upper panel: Minimum Antarctic total ozone for September to November, as observed by TOMS

and as computed by the participating models. For the transient models, the solid lines show the results of a

Gaussian smoothing applied to the individual year’s results (also plotted). For the timeslice experiments, the

mean and twice the standard deviation of the values within each sample are recorded by the colored triangles

and the error bars. For MAECHAM/CHEM only: (1) the values have been plotted 2 years late for clarity, and

(2) a standard tropospheric column of 40 DU has been added to the computed stratospheric columns. Lower

panel: The maximum area of the ozone hole during the period September to November, as given by the 220-

DU contour, for TOMS observations and participating models. The TOMS data and transient model results

are plotted for each year and with a Gaussian filter applied (solid lines). For the timeslice experiments the

means and twice the standard deviations of the model samples are included. For MAECHAM/CHEM only: (1)

the values have been plotted 2 years late for clarity, and (2) a standard tropospheric column of 40 DU has

been added to the computed stratospheric columns.
GISS model has a larger response than the other models,

with the simulation indicating a minimum in the decadally

averaged results of about 175 DU compared with almost

100 DU higher in the other transient runs. The date of

minimum Arctic ozone, again as indicated by the minimum of the decadally averaged results, varies from 2004

for the CCSR/NIES model to 2019 for the GISS model.

UMETRAC indicates a minimum at about the year 2015,

but the simulation ends shortly afterwards and the results

do not change substantially in the final decade. All three

transient runs indicate some delay in the onset of ozone

recovery, presumably due to increases in greenhouse gases

(GHGs), although such a result is subject to considerable

uncertainty because of the large interannual variability.

Although the timeslice experiments do not have the temporal resolution to give a precise indication of the timing

of future ozone recovery, the E39/C model results

(Schnadt et al., 2002) may go against the transient model

results by suggesting that increases in planetary waves

occur in the Arctic, speeding up ozone recovery. This may

be considered the “dynamical effect on chemistry”:

increases in planetary waves transport more ozone as well

as raise temperatures and decrease heterogeneous chemistry. Therefore, the net effect on ozone is that of the two

potentially competing processes of dynamics and radiation. If planetary waves increase, the “dynamical” effect

increases ozone and the “radiative” effect decreases

ozone, giving a relatively small response. If planetary

waves decrease, both the “dynamical” and “radiative”

effects are negative, leading to enhanced ozone depletion.

To resolve from the timeslice simulations whether

increases in GHGs are delaying the onset of ozone

recovery would require more results for the period 1990

to 2015.

In the Antarctic, the runs are all in fairly good

agreement. Of the transient runs, as in the Arctic, the

CCSR/NIES model indicates the earliest start of ozone

recovery (2001) followed by UMETRAC (2005) and

GISS (2008). The minima in the decadally averaged

results are all comparable (109, 86, and 98 DU, respectively). On the basis of the decadally averaged model

results this would appear to indicate that ozone recovery

would begin earlier in the Antarctic than in the Arctic.

Such an earlier start to recovery would also be detectable

earlier in observations in Antarctica, because of the

smaller interannual variability.

Observations of the size of the ozone hole (Figure

3-47; see also Figure 3-7) do not indicate any clear

recovery by October 2001, with interannual variability

now dominating over the current trends. For the GISS

model the maximum ozone hole is in 2015, whereas for

UMETRAC the maximum occurs in 2007, although any

systematic variation is obscured by the large interannual

variability noted previously. In contrast the CCSR/NIES

results show a peak as early as the year 2002.

3.5.3.3 THE 2020-2060 TIME FRAME: COMPLETE

OZONE RECOVERY

Those models that have run beyond the year 2020

indicate some recovery in ozone. Of particular importance is the return to “1980-like conditions,” when the

effects of anthropogenic halogen concentrations in the

polar regions are negligible. As noted in the previous

Assessment (Chapter 12 of WMO, 1999), this recovery

would be to a different vertical distribution of ozone, with

higher middle and upper stratospheric ozone due to the

change in the vertical temperature profile (see Chapter 4).

Using a 2-D model, Rosenfield et al. (2002) determined

the date for the recovery of total ozone to 1980 levels as a

function of day of year and latitude. In the Arctic, this

recovery was latest at the end of spring (after 2050) and

earliest in autumn (before 2035). Further, the impact of

CO2 increases was shown to accelerate the recovery from

that due to chemical changes alone by increasing the

downwelling over the poles. In contrast, if methane

amounts do not increase at the current rate, ozone recovery

could be slowed down in the future by the increased

importance of NOx chemistry (Randeniya et al., 2002).

Over Antarctica, downwelling is less important in

speeding up the ozone recovery.

Figures 3-46 and 3-47 show similar results for the

spring for 3-D models. However, in the Arctic, most models

do not show substantial ozone change throughout the

period 2020 to 2050, and the low values of the GISS model

for the decade 2010 to 2020 are no longer present after

2030 (Figure 3-46). In the Antarctic, the recovery of

spring ozone, already under way by 2020, continues in

the simulations completed (Figure 3-47). Recovery to

1980-like conditions occurs in the CCSR/NIES and GISS

models by about 2045, and perhaps a decade later in the

results of Austin et al. (2001). The CCSR/NIES and GISS

transient model results suggest a near-monotonic recovery

of ozone, but the results of Austin et al. (2001) suggest

that ozone could undergo further loss over the period 2025

to 2045. This was identified as due to increases in ice

PSCs as the lower stratospheric climate cools (but would

need to be confirmed by model simulations with more

detailed PSC schemes). Of the timeslice experiments, the

MAECHAM/CHEM results indicate a significant (but not

“full”) recovery in the Antarctic by 2030, consistent with

the transient experiments. In the Arctic this model also

simulates full or near-full recovery by 2030.

POLAR OZONE

3.79
3.5.3.4 ATTRIBUTION OF MODEL OZONE CHANGES

The contrast between the modeled Antarctic and

Arctic ozone behavior is striking in terms of the extent to

which the observed behavior can be regarded as a forced

response to chlorine loading. There is no question but

that this is the case in the Antarctic; the models, despite

their various imperfections, together give a clear, statistically significant signal that is in line with the observations. However, this is not the case in the Arctic: although

there is a general ozone decrease in the models during the

period of chlorine loading, it is not statistically significant. This is partly because of the weaker chemical impact

of chlorine loading in the Arctic than in the Antarctic, and

partly because of the greater natural variability in the

Arctic.

In a coupled chemistry-climate model, the attribution of ozone changes to dynamical and chemical

processes may be ambiguous because the dynamical

changes themselves may have been caused by chemical

changes to the ozone amounts. This is discussed further

in Chapter 4. Figure 3-45, illustrating the approximate

amounts of PSCs in the model simulations, should in principle reflect the amount of chemical ozone depletion (cf.

Figure 3-33). This would suggest, for example, that UMETRAC has slightly less Arctic ozone depletion relative to

observations, and this is reflected in the Arctic ozone trend

(Figure 3-46; Table 3-5) that is smaller than observed,

although the difference is not statistically significant.

Also, both MAECHAM/CHEM and E39/C results for

PSCs in Figure 3-45 suggest that their chemical ozone

depletion is larger than observed in the Arctic, although

the net ozone trend is similar to observations. The implication is that transport into the polar regions is enhanced

to compensate. However, this does not appear to be consistent with earlier results (e.g., Hein et al., 2001), which

if anything indicate reduced transport into the polar

regions.

In the Antarctic, many of the models have similar

ozone depletion rates, similar amounts of PSCs, and similar ozone trends, when allowance is made for the different

periods under consideration (Figure 3-47; Table 3-5). The

main exceptions are the GISS model for which PSC diagnostics are not available, and CMAM, which has less

ozone depletion than that observed in 2000. These inconsistencies in both hemispheres suggest the need for further investigation of the sizes of the transport versus chemical depletion terms in all the models included herein.

Considering now the future evolution of Arctic

ozone, the envelope of ozone variability defined by the

timeslice experiments in Figure 3-46b returns to higher

values in the future. Thus, the impact of the WMGHGs

on Arctic ozone is still uncertain. The two transient experiments performed for this Assessment, shown in Figure 3-

46a, are consistent with the timeslice experiments.

However, the transient experiment from the previous

Assessment falls well outside the envelope of variability

defined by the timeslice experiments over the coming several decades, particularly in terms of its lowest values.

The prediction of the previous Assessment that changes

in WMGHGs could lead to Arctic ozone minima of 100

DU over the coming several decades is not supported by

the more recent models. The reason for the difference

from the previous Assessment is that the more recent

models do not predict the severe decrease in planetary

wave drag over the Arctic (leading to essentially

Antarctic-like meteorological conditions in the Arctic in

some years) predicted by the earlier model. The earlier

caveats, including the need for ensemble integrations,

should again be borne in mind.

3.5.4 Summary

The main uncertainties of 3-D coupled chemistryclimate models stem from the performance of the underlying dynamical models. Temperature biases lead to errors

in the spatial extent of PSCs and the degree of chemical

ozone depletion. The model results also suggest significant differences in the transport of ozone to high latitudes,

although this is in need of further clarification. At the current stage of model performance, uncertainties in the

details of PSC formation and sedimentation are probably

less important in simulating ozone amounts than the

model temperature biases. Nonetheless, the accurate representation of PSC processes will prove to be increasingly

important as temperature biases become smaller by, for

example, the inclusion of non-orographic gravity wave

parameterizations. Another uncertainty is the amount of

aerosol present due to future unpredictable volcanic eruptions. For a large eruption such as that of Mt. Pinatubo,

sufficient aerosol would be present to provide additional

sites for heterogeneous chemistry and possible severe

ozone loss for a period of a few years (see, e.g., Tabazadeh

et al., 2002), although this perturbation would not affect

the long-term ozone trend. The impact of volcanic eruptions on coupled chemistry-climate model results has not

been discussed in this section, but further details may be

found in Chapter 4.

For the transient model simulations, the start of

ozone recovery, as defined in Section 3.5.3, occurs in the

Antarctic within the next decade, depending on the model,

and in the Arctic it occurs within the next two decades. In

the Antarctic, however, model results suggest that the vertical and horizontal extent of the ozone hole may increase

slightly further over the next few years. Thus, the results

here suggest that the start of ozone recovery will occur

POLAR OZONE

3.80
slightly later in the Arctic than in the Antarctic. Further,

since the halogen amounts are thought to have maximized

in 2001 (see Chapter 1, Figure 1-7), the start of ozone

recovery in the Arctic in the models is delayed by up to

almost two decades, probably by greenhouse gas

increases. Most of the models come to similar conclusions on this issue, but one of the coupled chemistryclimate model experiments (Schnadt et al., 2002) suggests

that greenhouse gas increases would tend to speed up

rather than slow down ozone recovery in the Arctic. It

should also be recognized that interannual variability on

the sub-decadal time scale may still lead to ozone

extremes. In the worst-case scenario, therefore, it may

take until at least the end of the 2020s before we can be

certain that ozone recovery has started in the Arctic. To

put this into perspective, most models predict relatively

modest changes in future spring Arctic column ozone

(under 10%). The one model that does predict a major

Arctic ozone change in the near future (described in

Shindell et al., 1998a) has lower spatial resolution, simplified ozone transport, and parameterized ozone

chemistry. In comparison, the other models have more

accurate treatments of these processes.

On the longer time scale, to the middle of the 21st

century, model predictions appear to be more uncertain.

Hence, although recovery of Antarctic ozone to 1980-like

conditions (“full ozone recovery”) is to be expected by

about 2050, models will need to have a better representation of the water vapor increase than has hitherto been

possible, as well as an accurate specification of methane

changes, for full confidence in their predictions. Although

the results for the Arctic are less certain, since most models

indicate relatively modest change in ozone, it is possible

that “full recovery” may occur somewhat earlier there.

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Appendix 3A

SATELLITE MEASUREMENTS IN THE ANTARCTIC AND ARCTIC

In this appendix, we update information on satellite instruments that measure ozone and other species that are pertinent to polar ozone issues. We briefly review the status of a variety of ozone and related measurements for the Antarctic

and Arctic. Total ozone observations and ozonesondes have been extensively discussed in previous reports, and are further reviewed in Chapter 4 of this Assessment.

Earth Probe (EP) Total Ozone Mapping Spectrometer (TOMS): Data from the TOMS instrument have been

extensively used to track Arctic and Antarctic ozone changes. The TOMS data are discussed in Chapter 4, Appendix 4A.

The EP/TOMS operational processing configuration has recently been changed in order to apply a correction to a crosstrack bias error that has grown since 2000. Although 2001 data are included herein, they are of slightly greater uncertainty (see discussion in the Earth Probes TOMS description in Chapter 4, Appendix 4A).

Microwave Limb Sounder (MLS): MLS observations started in September 1991 with the launch of the Upper

Atmosphere Research Satellite (UARS). It gives column and profile data on ozone, ClO, and HNO3. Recently, the satellite and instrument have experienced problems, and the instrument is turned on only when there are situations of particular interest. In particular, MLS was taken out of standby mode on 31 January 2000 and was operated during the 2-13

February and 27-29 March periods in conjunction with the SOLVE/THESEO 2000 campaign (Santee et al., 2000).

Halogen Occultation Experiment (HALOE): HALOE observations also started in September 1991 with the

launch of UARS. HALOE is a solar occultation instrument that makes measurements in the IR at both sunset and sunrise.

HALOE measurements are used to retrieve profiles of ozone, HF, HCl, CH4, H2O, NO, NO2, and aerosol extinction.

HALOE occultation latitudes are variable over the course of the northern winter, but do not reach the high northern latitudes. During the winter of 1999/2000, the maximum latitude sampled by HALOE was approximately 63°N in midMarch. HALOE typically has difficulty sampling the polar vortex during midwinter, but does sample the vortex edge

region in the fall (Pierce et al., 2002).

Global Ozone Monitoring Experiment (GOME): The GOME instrument was launched in 1995 aboard the

second European Remote Sensing (ERS-2) satellite. It measures column ozone, NO2, BrO, and OClO, and ozone profiles. Details on the GOME instrument can be found in Burrows et al. (1999). The GOME data have been subject to validation exercises (Hansen et al., 1999; Hoogen et al., 1999; Corlett and Monks, 2001; Ionov et al., 2001; Piters et al., 1999;

Rathman et al., 1997).

Polar Ozone and Aerosol Measurement (POAM): POAM is a solar occultation instrument that provides ozone,

water vapor, NO2, and aerosol extinction profiles in the polar regions. POAM II was launched in 1993 aboard the French

Satellite Pour l’Observation de la Terre (SPOT)-3 satellite. These measurements were interrupted by the failure of the

SPOT-3 satellite in November 1996. POAM III was subsequently launched on the French SPOT-4 satellite in March

1998 and is currently operational.

Lucke at al. (1999), Lumpe et al. (2002), and Randall et al. (2002) have published intercomparisons between

POAM III and other instruments. Results from POAM III include studies on dehydration (Nedoluha et al., 2000), ozone

loss (Randall et al., 2002; Hoppel et al., 2002), and PSCs (Bevilacqua et al., 2002).

Improved Limb Atmospheric Spectrometer (ILAS): ILAS is a satellite instrument that uses the solar occultation technique (Sasano et al., 1999; Nakajima et al., 2002). ILAS was launched onboard the ADvanced Earth Observing

Satellite (ADEOS) on 17 August 1996. ILAS made measurements over high-latitude regions from 57°N to 71°N and

from 64°S to 88°S from late October 1996 until late June 1997.

ILAS consists of an infrared spectrometer that covers the wavelength region from about 6 to 12 mm. ILAS made

vertical profile measurements of ozone, HNO3, NO2, N2O, CH4, and H2O from the infrared spectrometer, as well as vertical profiles of aerosol extinction coefficient at 780 nm from the visible spectrometer. Ozone data are validated by correlative and coincident measurements from several instrumental techniques (Sugita et al., 2002). Nitric acid, NO2, and

H2O data have also been validated (Koike et al., 2000; Kanzawa et al., 2002; Irie et al., 2002).
CHAPTER 4

Global Ozone: Past and Future

Contributors:

J. Austin

M. Bourqui

P. Braesicke

C. Brühl

N. Butchart

D.B. Considine

D. Cunnold

D.W. Fahey

E.L. Fleming

P.M. de F. Forster

M.A. Geller

S. Godin-Beekmann

V. Grewe

J.D. Haigh

C.H. Jackman

P. Johnston

U. Langematz

K.S. Law

J.B. Liley

I.A. Megretskaia

A.J. Miller

T. Nagashima

G. Pitari

R.W. Portmann

V. Ramaswamy

B. Rognerud

J.E. Rosenfield

M.N. Ross

C. Schnadt

M.D. Schwarzkopf

D.T. Shindell

C.A. Smith

S. Smyshlyaev

R.S. Stolarski

G.J.M. Velders

R. Wang

D.K. Weisenstein

F. Wu

R.J. Zander

Lead Authors:

M.P. Chipperfield

W.J. Randel

Coauthors:

G.E. Bodeker

M. Dameris

V.E. Fioletov

R.R. Friedl

N.R.P. Harris

J.A. Logan

R.D. McPeters

N.J. Muthama

T. Peter

T.G. Shepherd

K.P. Shine

S. Solomon

L.W. Thomason

J.M. Zawodny

CHAPTER 4

GLOBAL OZONE: PAST AND FUTURE

Contents

SCIENTIFIC SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1

4.1 INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.5

4.2 DESCRIPTION OF PAST CHANGES IN OZONE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.6

4.2.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.6

4.2.2 Changes in Total Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.7

4.2.2.1 Global Ozone: NH and SH Midlatitudes and Tropics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.7

4.2.2.2 Seasonal Changes in Total Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.10

4.2.3 Changes in the Vertical Profile of Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.16

4.2.3.1 Upper Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.16

4.2.3.2 Lower Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.18

4.2.4 Vertical Integral of the Profile Trends vs. Column Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.19

4.2.5 Tropospheric Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.22

4.2.6 Solar Cycle Effects on Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.22

4.2.6.1 Column Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.22

4.2.6.2 Vertical Profile of the Solar Signal . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.24

4.2.6.3 Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.25

4.3 CURRENT UNDERSTANDING OF PAST CHANGES IN STRATOSPHERIC AEROSOL,

WATER VAPOR, AND NO2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.25

4.3.1 Stratospheric Aerosol Variability . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.25

4.3.2 Stratospheric Water Vapor Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.27

4.3.3 Trends in Stratospheric NO2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.28

4.4 CURRENT UNDERSTANDING OF PAST CHANGES IN STRATOSPHERIC TEMPERATURE . . . . . . . 4.29

4.4.1 Updated Observations of Stratospheric Temperature . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.29

4.4.2 Causes of Past Temperature Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.30

4.4.2.1 Simulated Temperature Trends Due to Observed Ozone Trends . . . . . . . . . . . . . . . . . . . . . 4.31

4.4.2.2 Simulated Temperature Trends Due to Increased WMGG Concentrations . . . . . . . . . . . . . 4.33

4.4.2.3 The Role of Trends in Stratospheric Water Vapor . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.34

4.4.2.4 Synopsis of Simulated Temperature Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.35

4.5 CHEMICAL INFLUENCE ON PAST CHANGES IN OZONE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.37

4.5.1 Chemical Processes that May Contribute to Ozone Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.37

4.5.1.1 Gas-Phase Chemical Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.37

4.5.1.2 Midlatitude Lower Stratospheric Heterogeneous Processes . . . . . . . . . . . . . . . . . . . . . . . . 4.38

4.5.2 Quantifying the Chemical Influence on Ozone: Modeling Tools . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.41

4.5.2.1 2-D Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.41

4.5.2.2 3-D Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.42

4.5.3 Quantifying the Chemical Influence on Ozone: Model Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.43

4.5.3.1 Effects of NOx- and ClOx-Related Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.43

4.5.3.2 Effect of Polar Vortex on Midlatitudes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.45

4.5.3.3 Effect of Water Vapor Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.45

4.5.3.4 Assessment Model Simulations of Past Ozone Changes . . . . . . . . . . . . . . . . . . . . . . . . . . 4.46
4.6 DYNAMICAL INFLUENCE ON PAST CHANGES IN OZONE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.52

4.6.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.52

4.6.2 Dynamical Processes that May Contribute to Ozone Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.54

4.6.2.1 Changes in Planetary-Wave Drag . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.54

4.6.2.2 Effects of Tropospheric Circulation and Tropopause Changes . . . . . . . . . . . . . . . . . . . . . . 4.55

4.6.2.3 Hemispheric Variations and Annular Modes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.56

4.6.3 Statistical Analysis of Observations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.56

4.6.3.1 Changes in Planetary-Wave Driving . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.56

4.6.3.2 Changes in Tropospheric Circulation and Tropopause Height . . . . . . . . . . . . . . . . . . . . . . 4.57

4.6.3.3 Coupling Issues . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.58

4.6.4 Modeling Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .4.58

4.6.5 Cause of the Dynamical Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.60

4.6.6 Effects of Mt. Pinatubo Eruption . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.61

4.6.7 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.62

4.7 IMPLICATIONS FOR CLIMATE OF UPDATED STRATOSPHERIC OZONE CHANGES . . . . . . . . . . . . 4.62

4.8 FUTURE CHANGES IN OZONE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.63

4.8.1 Effect of Decreasing Halogens . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.64

4.8.2 Influence of Changing Climate Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.65

4.8.3 Other Stratospheric Influences . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.66

4.8.3.1 Dynamical Effects . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.66

4.8.3.2 Future Effects of Aircraft . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.67

4.8.3.3 Rocket Emissions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.67

4.8.4 Future Tropospheric Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.68

4.9 SYNTHESIS OF CURRENT UNDERSTANDING OF PAST AND FUTURE CHANGES IN OZONE . . . . 4.70

4.9.1 Discussion of Past Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.70

4.9.2 Discussion of Future Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.72

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.72

APPENDIX 4A: DESCRIPTION OF OZONE DATASETS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.85

4A.1 Ground-Based Total Ozone Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.85

4A.2 Satellite Total Ozone Measurements and Merged Datasets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.85

4A.3 Ozone Profile Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.88

APPENDIX 4B: 2-D MODEL SCENARIOS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.90
SCIENTIFIC SUMMARY

Total Column Ozone

• Global mean total column ozone for the period 1997-2001 was approximately 3% below the pre-1980 average

values. Since systematic global observations began in the mid-1960s, the lowest annually averaged global total

column ozone occurred in 1992-1993 (5% below the pre-1980 average). These changes are evident in each available global dataset.

• No significant trends in total column ozone have been observed in the tropics (25°N-25°S) for 1980-2000. A

decadal variation of total column ozone (with peak-to-trough variations of ~3%) is observed in this region, approximately in phase with the 11-year solar cycle. Total column ozone trends become statistically significant in the latitude bands 25°-35° in each hemisphere.

• There are a number of differences in total column ozone behavior between the two hemispheres:

• Averaged for the period 1997-2001, total column ozone in the Northern Hemisphere (NH) and Southern

Hemisphere (SH) midlatitudes (35°-60°) are about 3% and 6%, respectively, below their pre-1980 average

values.

• The seasonality of total column ozone changes (1997-2001 relative to pre-1980) is different in the NH and SH

midlatitudes. Over NH midlatitudes, larger ozone decreases are observed during winter-spring (~4%); summerautumn decreases are approximately half as large. Over SH midlatitudes, long-term ozone decreases exhibit

similar magnitude (~6%) during all seasons.

• Pronounced negative anomalies are observed in the NH midlatitude time series during 1992-1995 in the winterspring seasons. Similar anomalies are not seen in the SH midlatitudes.

• There is a sharp drop in ozone at SH midlatitudes during 1985-1986. A similar drop is not observed in the NH.

Vertical Ozone Distribution

• Ozone profile trends derived from Stratospheric Aerosol and Gas Experiment (SAGE I and SAGE II, version v6.1)

data show significant negative trends over latitudes 60°N to 60°S for altitudes ~35-50 km (with extremes near 40

km). Trend maxima of 7-8%/decade over the period 1979-2000 are observed in latitude bands 35°-60° of both

hemispheres, with no significant interhemispheric difference. These satellite results are in good agreement with

independent Umkehr ozone measurements available over the NH midlatitudes. Ozone decreases in the upper stratosphere (above 25 km) contribute 15-30% to changes in the total column at midlatitudes.

• The updated SAGE data reveal significant negative trends in the small amount of ozone above 30 km extending

throughout the tropics, a feature not observed in previous Assessments based on shorter time records.

• Ozonesonde measurements over the NH midlatitudes show that ozone between 20 and 27 km decreased continuously during 1980-2000, while ozone between 10 and 20 km decreased through the early 1990s and was relatively

constant thereafter. This behavior is consistent with observed changes in NH midlatitude column ozone.

• The ozonesonde measurements show no trend in tropospheric ozone for 1980 to 2000 for most stations, and the

mean trend averaged over all midlatitude stations is also zero.

Ozone-Related Constituents

• Stratospheric aerosol variability over the past 25 years has been dominated by the effects of episodic volcanic eruptions, with subsequent recovery. Following the large eruption of Mt. Pinatubo in 1991, relaxation to a nonvolcanic

level continued to at least 1999. There is currently no evidence of a trend in the nonvolcanic aerosol loading over

the past 25 years.

GLOBAL OZONE

4.1
• Stratospheric water vapor measurements at Boulder, Colorado, U.S. (40°N), for the period 1981-2000 show a statistically significant increase of approximately 1%/year over altitudes 15-28 km. For the shorter period 1991-2001,

global satellite measurements covering latitudes 60°S to 60°N show a similar trend of 0.6-0.8%/year for altitudes

~25-50 km, but no significant trend at lower altitudes. This increase in water vapor is substantially larger than can

be explained by tropospheric methane (CH4) trends. Characterization of stratospheric water vapor trends is limited

by the lack of global long-term measurements.

• Stratospheric column nitrogen dioxide (NO2) measurements from Lauder, New Zealand (45°S), for the period

1981-2001, and Jungfraujoch, Switzerland (46°N), for 1985-2001, show statistically significant positive trends of

approximately 6%/decade. Transient decreases are also observed following the El Chichón and Mt. Pinatubo volcanic eruptions, which are broadly simulated by models that include heterogeneous chemistry on sulfate aerosols.

Stratospheric Temperature Trends

• Satellite observations indicate that, on an annual and global mean, the stratosphere has cooled over the last two

decades. In the lower stratosphere, global and annual mean temperatures for the late 1990s are approximately 1 K

lower than values in the late 1970s. Significant annual-mean cooling of the lower stratosphere for the past two

decades is found over midlatitudes of both hemispheres (approximately 0.6 K/decade), but no significant trends are

observed near the equator. The annual mean temperature trends in the upper stratosphere are larger, with an approximately globally uniform cooling for 1979-1998 of about 2 K/decade near the stratopause (~50 km).

• Modeling studies indicate that changes in ozone, well-mixed greenhouse gases, and stratospheric water vapor can

explain the major features of the observed global and annual-mean stratospheric cooling over the past two decades.

Cooling due to ozone depletion dominates over the impact of well-mixed greenhouse gases in the lower stratosphere. Upper stratospheric temperature trends are due, roughly equally, to ozone and well-mixed greenhouse gas

changes. At 20 km, the observed annual-mean cooling in northern midlatitudes is much larger than that simulated

using observed ozone and greenhouse gas changes. In this region increases in water vapor may have been as important as ozone loss in determining the cooling. The inclusion of water vapor generally improves agreement with

observations, but uncertainties in the spatial and temporal distribution of its change are large.

Attribution of Past Changes in Ozone

• The vertical, latitudinal, and seasonal characteristics of changes in midlatitude ozone are broadly consistent with

the understanding that halogens are the primary cause of these changes, in line with similar conclusions from the

1998 Assessment.

• Assessment models forced by observed changes in halocarbons, source gases, and aerosols broadly reproduce the

long-term changes observed in midlatitude total column ozone (35°N-60°N and 35°S-60°S) for 1980-2000, within

the uncertainties of the observations and model range. However, the range of model results is large over SH midlatitudes, which is at least partly due to their treatment of the Antarctic ozone hole. In addition, models suggest that

the chemical signal of ozone loss following the major eruption of the Mt. Pinatubo volcano in the early 1990s

should have been symmetric between hemispheres, but observations show a large degree of interhemispheric asymmetry in midlatitudes.

• There is increased evidence that changes in atmospheric dynamics have had a significant influence on NH midlatitude column ozone on decadal time scales. Natural variability, changes in greenhouse gases, and changes in column

ozone itself are all likely to contribute to these dynamical changes. Furthermore, because chemical and dynamical

processes are coupled, their contributions to ozone changes cannot be assessed in isolation.

• The observed ozone depletion in the upper stratosphere is consistent with observed changes in anthropogenic chlorine.

The vertical and latitudinal profiles of trends in the upper stratosphere are reproduced by photochemical models, but

the magnitudes of changes are sensitive to concurrent trends in temperature and other constituents such as CH4.

GLOBAL OZONE

4.2
• Stratospheric planetary-wave drag drives the wintertime ozone buildup in the extratropics, so the observed decrease

in NH planetary-wave drag likely has contributed to the observed decrease in NH total column ozone over the last 20

years. The relationship between planetary-wave drag and ozone is understood, but its quantification in observations

is relatively crude.

• There is an observed relationship between total column ozone and tropopause height, and several tropospheric circulation indices. These indices have changed over the last 20 years in the NH in such a way as to imply a decrease in

total column ozone. However, the causality of these relationships is not well understood, which makes their extrapolation to long time periods problematic.

Future Changes in Ozone

• The expected decrease in stratospheric chlorine loading over the next 50 years is predicted to lead to an increase in

the global total column ozone, although there are differences in the rate of increase between different twodimensional assessment models. Future ozone levels will also be influenced by other changes in atmospheric composition and by climate change. Because of year-to-year variability it could take as long as a decade to demonstrate a

leveling of total column ozone.

• Stratospheric cooling (due mainly to projected carbon dioxide (CO2) increases) is predicted to enhance future ozone

increases in the upper stratosphere. This effect in isolation would substantially hasten the predicted rate of total

ozone increase in coming decades. However, a reliable assessment of effects on total column ozone is limited by

uncertainties in the lower stratospheric response to these changes.

• Projected increases in methane (CH4) and nitrous oxide (N2O) are predicted to have small effects on the rate of

increase of global column ozone in the next 50 years, when chlorine changes are the dominant effect. After that time,

changes in CH4 and N2O become relatively more important.

• Future large volcanic eruptions would be expected to cause transient enhancements of chemical ozone depletion

(lasting 3-4 years) but should not alter the overall predicted evolution of ozone.

• Future changes in tropospheric ozone are highly dependent upon the scenario adopted for future emissions of ozone

precursors, but all scenarios adopted by the Intergovernmental Panel on Climate Change (IPCC) lead to increases in

tropospheric ozone up to 2050. Allowing for climate change leads to smaller increases in tropospheric ozone compared with simulations using present-day climate.

GLOBAL OZONE

4.3

GLOBAL OZONE

4.5

4.1 INTRODUCTION

Concentrations of ozone in the atmosphere can

change as a result of both natural processes and human

activities, and accurate knowledge of past changes is

important both for attribution of causes and for prediction

of future changes. Past ozone assessments have documented long-term decreases in stratospheric ozone and

have devoted separate chapters to ozone observations,

interpretation of past changes, and predictions of future

change. In contrast, this chapter takes a comprehensive

view of past and future changes in global ozone, combining observational and modeling results. In brief, we

seek to address the following questions:

• What have been the changes to stratospheric ozone

in the past, and how well do we understand those

changes?

• What is our best estimate for stratospheric ozone in

the future, and what are the main uncertainties in this

estimate?

The bulk of atmospheric ozone is found in the stratosphere, with approximately 90% of the vertically integrated column amount occurring above the tropopause

(the “ozone layer”). Long-term changes of stratospheric

ozone are related to changes in tropospheric source gases

(transported into the stratosphere), to variations in solar

radiation and volcanic effects, and to changes in largescale transport. Changes in tropospheric ozone (which is

an environmental pollutant near the surface) are mainly

associated with increased urban pollution and biomass

burning. Stratospheric ozone has decreased over the last

several decades, while tropospheric ozone has increased

(at least over Northern Hemisphere (NH) midlatitudes,

where long-term observations are available; Logan et al.,

1999). In the extratropics the stratospheric decreases

greatly outweigh the tropospheric increases in terms of

column amount, and therefore observed column ozone

changes primarily reflect the stratospheric decreases.

Long-term changes in stratospheric ozone are the primary

focus of this chapter, and tropospheric ozone is discussed

only briefly (in regard to its contribution to past and future

column ozone changes). Detailed discussions of tropospheric ozone are included in IPCC (2001); some of those

results are highlighted here where relevant.

The distribution of ozone in the atmosphere is

maintained by a balance between photochemical production and loss, and by transport between regions of net production and net loss. Understanding long-term changes

in ozone, both past and future, requires knowledge of these

processes and a quantification of their relative importance

and coupling. The basic structure and zonal mean circulation of the stratosphere are summarized in Figure 4-1.

There is an important distinction between the different

vertical regions in the stratosphere, which can be characterized by their radiative, dynamical, and transport

behavior. The photochemical lifetime of odd-oxygen

(Ox = O + O3, where O is atomic oxygen and O3 is ozone)

varies strongly with altitude between these regions (and

also with latitude and season; e.g., Garcia and Solomon

(1983)). (In the stratosphere, O and O3 are in a rapidly

established equilibrium, and so any analysis of processes

that control O3 should be based on the family Ox.) The

lowermost stratosphere (LMS) covers the region between

the tropopause and the isentropic surface of about 380 K

(about 15 km), and this region is distinct because air can

be transported between the troposphere and stratosphere

by isentropic transport alone. Above the LMS, between

the 380-K surface and about 25 km is the lower stratosphere (LS). In both the LMS and LS, the lifetime of oddoxygen is long (~months-years), and odd-oxygen will

react to chemical changes slowly and be strongly influenced by atmospheric transport. The region between 30

and 50 km is typically referred to as the upper stratosphere

(US) based on radiative/dynamical behavior. However,

the odd-oxygen photochemical lifetime is relatively short

(less than transport time scales) above 25 km, hence dominated by photochemical control, and for simplicity we

refer to the region from 25 to 50 km as the US throughout

this chapter.

By considering the factors that control the abundance of ozone in the stratosphere, we can classify the

three different regions in Figure 4-1 by the complexity of

the processes that need to be understood. The US can be

Figure 4-1. Schematic diagram showing the structure of, and transport within, the stratosphere. The

three stratospheric regions (LMS, LS, US) discussed in the text are indicated.
GLOBAL OZONE

4.6

regarded as the easiest region to analyze, because ozone

there is under direct photochemical control and direct

transport effects can be neglected (although transport

affects the abundance of ozone-destroying species). The

LS and LMS are more difficult to analyze, because both

transport and chemistry directly affect ozone. Within

these lower altitude regions, the LMS region can be considered more difficult than the LS because it is directly

coupled by isentropic transport to the troposphere.

As discussed in the previous Assessment (WMO,

1999), there is strong evidence that the main driver for

the upper stratospheric ozone loss near 40 km and the

polar lower stratospheric springtime loss is halogen chemistry. In the upper stratosphere, ozone loss occurs through

purely gas-phase chlorine chemistry. In the polar lower

stratosphere, heterogeneous chemistry in or on polar stratospheric clouds (PSCs) causes the conversion of stable

chlorine reservoirs to more active forms. Ozone loss then

occurs through reactions involving chlorine and bromine

species (see Chapter 3). It is also well established that

some of the chemical depletion that occurs during spring

in the high latitudes must eventually lead to a decrease

(or “dilution”) of ozone at midlatitudes when the polar

vortices break down. However, this process is poorly

quantified at present. A number of other processes, both

chemical and dynamical, have been proposed as contributing to ozone decreases in the midlatitude lower stratosphere. The chemical processes include wintertime

transport of chlorine-activated air from the vortex followed by chemical loss, and in situ loss initiated by heterogeneous chemistry on liquid sulfate aerosols. The

dynamical processes include changes in planetary-wave

driving (and associated stratospheric transport circulation), and changes in tropospheric circulation and in

tropopause height. While there is good mechanistic

evidence that these processes occur and affect ozone,

quantifying their contribution to the long-term change is

difficult. Furthermore, the chemical and dynamical contributions are coupled and not truly separable, but rather

act synergistically.

Past ozone assessments have focused on estimates

of linear trends in ozone, usually determined via a multiple regression analysis (including terms to account for

other known sources of periodic variability, such as the

solar cycle or quasi-biennial oscillation). The choice of

fitting a linear trend was based on the expected response

to the approximate linear increases in ozone-depleting

substances (ODSs) such as chlorofluorocarbons (CFCs)

and halons (prior to the middle 1990s). However, this

linear increase in ODS has not continued beyond the

middle 1990s (see Chapter 1), and the continued representation of ozone variations as simple linear trends is not

appropriate. The philosophy adopted here is to focus primarily on the observed time series of ozone variations

updated through 2001, including their detailed latitude

and altitude structure, and use these time variations as fingerprints of the mechanisms that influence change. Linear

trends are calculated for some variables, namely, for ozone

profile variations (where data availability limits the information in time series), and in comparison of the model

results to ozone observations in Section 4.5 (where trends

are a useful summary diagnostic). For clarity, we reserve

the use of the term “trend” to refer to linear trends calculated over a specific time period (typically determined by

regression fit).

The overall layout of this chapter is as follows.

Section 4.2 provides an update of observed changes in

ozone through 2001, including the latitudinal and temporal changes in column and profile ozone. We focus

on global or near-global variability (90°N-90°S or 60°N60°S), together with specific analyses of the tropics

(25°N-25°S) and Northern Hemisphere/Southern

Hemisphere midlatitudes (which are defined, somewhat

arbitrarily, as latitudes 35°N-60°N and 35°S-60°S). The

latitude bands 25°-35°N and S are transition zones,

where observations show that decadal ozone trends start

to become significant. Tropospheric ozone is briefly discussed, and its role in column ozone is estimated. We

also specifically discuss the 11-year solar cycle in ozone

evaluated from observations and models; this topic is

important for quantifying decadal variability in the relatively short observational record. Sections 4.3 and 4.4

describe the complementary observations of past

changes in stratospheric aerosol, water vapor, nitrogen

dioxide (NO2), and temperature, and summarize our

understanding of their causes. Section 4.5 describes our

current understanding of the chemical influence on

ozone changes at midlatitudes, and Section 4.6 discusses

the dynamical influences on midlatitude ozone changes.

Section 4.7 discusses what effect the observed changes

in ozone may have had on climate, and Section 4.8 provides updated estimates of how global ozone may evolve

in the future. An overall synthesis of current understanding of past and future global ozone changes is presented in Section 4.9. 

4.2 DESCRIPTION OF PAST CHANGES IN

OZONE

4.2.1 Introduction

Global and hemispheric-scale variations in stratospheric ozone can be quantified from extensive observational records covering the past 20-30 years. The goal
GLOBAL OZONE

4.7

here is to provide an updated description of interannual

variability in ozone over the most recent decades (approximately 1970-2000), with a focus on changes outside of

the polar regions (over latitudes approximately 60°N60°S). There are numerous ways to measure ozone in the

atmosphere, but they fall broadly into two categories:

measurements of total or column ozone (the vertically

integrated amount of ozone above the surface) and measurements of the vertical profile of ozone. There are more

independent datasets, longer time series, and better global

coverage for total ozone. Regular measurements of total

ozone are available from a network of surface stations,

mostly in the midlatitude Northern Hemisphere, with reasonable coverage extending back to the 1960s (although

geographical coverage is limited before the 1970s). Nearglobal, continuous total ozone data are available from

satellite measurements beginning in 1979. However,

changes in availability and calibration of ground-based

measurements and satellite datasets, and interruptions in

observational records, produce datasets in which the systematic errors are complicated functions of time.

Systematic errors of 1% are typical even for the most reliable instruments, whereas interannual changes in global

or hemispherically averaged ozone are typically a few

percent. The methodology adopted in the past (and used

here) is to compare estimates of ozone change from several different and independent techniques (i.e., groundbased and satellite datasets), and use their differences to

provide a measure of overall uncertainty. The results indicate an overall good agreement among the different data

sources for changes in total ozone, and thus we have reasonable confidence in describing the spatial and temporal

characteristics of past changes.

There are fewer independent datasets describing

long-term changes in the vertical profile of ozone.

Satellite observations are the only dataset with global profile coverage, but they are limited in vertical extent and

temporal coverage. Reasonably long time records are

available from balloonborne ozonesonde and groundbased Umkehr data, but these are primarily concentrated

in NH midlatitudes. Thus, in the case of ozone profile

measurements, direct comparison of long records from

independent datasets is possible only over NH midlatitudes. Over that region there is reasonably good agreement for long-term changes between the satellite and

Umkehr data (in the upper stratosphere), and satellite and

ozonesondes (in the lower stratosphere), and this suggests

that the satellite data are reliable for global analyses.

There is approximate but not exact agreement between

trends in NH midlatitude total ozone and the vertically

integrated profile changes (from ozonesondes in the lower

stratosphere, and satellite/Umkehr in the upper stratosphere); the lack of more detailed agreement is probably

related to the sparse ozonesonde sampling.

4.2.2 Changes in Total Ozone

4.2.2.1 GLOBAL OZONE: NH AND SH MIDLATITUDES

AND TROPICS

Five datasets of zonal and monthly mean total

ozone values developed by different scientific teams are

examined in this Assessment. These consist of various

combinations of ground-based and merged satellite

datasets, as described in Fioletov et al. (2002) and summarized in Appendix 4A (p. 4.86). The global and latitudinal structure of total ozone variability is analyzed based

on seasonal (3-month average) and zonally averaged time

series (on a 5° latitude grid). The ground-based dataset

extends in time over ~1965-2001 (depending on latitude,

with longer and more reliable records available in the midlatitude NH), and the satellite datasets span 1979-2001.

Details of the data analyses are discussed in Fioletov et

al. (2002). Briefly, the seasonal cycle of total ozone for

each dataset is removed by taking monthly averages for

each month of the year for the period 1979-1987, and subtracting these from the original data. Area-weighted ozone

deviations are then calculated and expressed as anomalies

with respect to the period 1964-1980. For regions with

missing data, ozone deviations are assumed identical to

the surrounding latitude belt where data are available,

prior to area averaging. Systematic differences between

the separate datasets of up to 3% are found (for global

average ozone), although the deseasonalized total ozone

deviations agree to within 0.9% (95% confidence).

Overall, each of the total ozone datasets gives similar estimates of large-scale changes; each of the datasets is

included in the figures below, and the spread between

them gives an estimate of the uncertainty due to measurement technique and sampling. 

The total ozone deviations for the 60°S-60°N latitude belt are shown in Figure 4-2 (top), and corresponding results for the global mean (90°S-90°N) are shown

in Figure 4-2 (bottom). All five datasets indicate very

similar ozone variations, with differences typically less

than 0.5%. The global ozone (60°S-60°N or 90°S-90°N)

amount shows overall decreasing values between the late

1970s to early 1990s, a relative minimum during 1992-

1994, and a slight increase during the late 1990s. A relatively large decadal variation of global total ozone is seen

in Figure 4-2, which is likely associated with the 11-year

solar cycle (e.g., Chandra and McPeters, 1994; Bojkov

and Fioletov, 1995; see further discussion in Section

4.2.6). Figure 4-3 shows the deseasonalized global anom-
GLOBAL OZONE

4.8

Figure 4-2. Deseasonalized, area-weighted seasonal (3-month average) total ozone deviations, estimated

from five different global datasets. Each dataset was deseasonalized with respect to the period 1979-1987,

and deviations are expressed as percentages of the ground-based time average for the period 1964-1980.

Results are shown for the region 60°S-60°N (top) and the entire globe (90°S-90°N) (bottom). The different

satellite datasets cover 1979-2001, and the ground-based data extend back to 1964. TOMS, Total Ozone

Mapping Spectrometer; SBUV, Solar Backscatter Ultraviolet; NIWA, National Institute of Water and

Atmospheric Research (New Zealand). Adapted from Fioletov et al. (2002).

1965 1970 1975 1980 1985 1990 1995 2000

-6

-4

-2

0

2

Ground-based data

TOMS zonal means

SBUV-SBUV/2

Merged satellite data

NIWA assimilated data

1965 1970 1975 1980 1985 1990 1995 2000

- 6

- 4

- 2

0

2

Total Ozone Adjusted for Seasonal Effects

60°S-60°N

90°S-90°N Deviation (%) Deviation (%)
GLOBAL OZONE

4.9

1965 1970 1975 1980 1985 1990 1995 2000

-6

-4

-2

0

2 Deviation (%)

G round-based data

TOMS zonal means

SBUV-SBUV/2

Merged satellite data

NIWA assimilated data

1965 1970 1975 1980 1985 1990 1995 2000

-6

-4

-2

0

2 Deviation (%)

Total Ozone Adjusted for Seasonal, QBO, and Solar Effects 

60°S-60°N

90°S-90°N

Figure 4-3. Deseasonalized, area-weighted total ozone deviations from five datasets (as in Figure 4-2),

adjusted for solar and QBO effects, for the region 60°S-60°N (top) and the entire globe (90°S-90°N) (bottom).

The solar and QBO effects were estimated using the ground-based data only and subtracted from all five

datasets. Adapted from Fioletov et al. (2002).
alies, after statistical removal of the solar cycle and quasibiennial oscillation (QBO)-related variations (based on

standard regression analyses, as discussed in detail in

SPARC (1998)). The details of the solar cycle statistical

fit, and possible aliasing of solar and volcanic signals in

ozone, are discussed in detail in Section 4.2.6. The time

series in Figure 4-3 show that the residual global ozone

variability can be described by an approximately linear

decline from the late 1970s to the early 1990s, a minimum

during 1992-1994, increases during the late 1990s, and

relatively low values in 2000-2001. The global ozone

levels for the years 1998-1999 are similar to those of the

early 1980s, but about 2% lower than the late 1970s

values. However, the relatively low global ozone values

in 2000-2001 suggest caution in interpreting the higher

values for 1998-1999. A smoothed version of the global

ozone anomalies is shown in Figure 4-4, based on

applying four passes of a 13-month running mean to the

global anomalies in Figure 4-3. These smoothed time

series highlight the low-frequency interannual changes,

and provide concise estimates of global ozone changes

for comparison with the assessment model results in

Section 4.5.

About 40% of global ozone is located in the tropics

between 25°N-25 °S, and variability in this region over

the period 1979-2001 is shown in the middle curves of

Figure 4-5. There is a strong decadal variation of total

ozone in the tropics, with peak-to-trough variations of

approximately 3%, and maxima approximately in phase

with the 11-year solar cycle (top panel in Figure 4-5).

After statistical removal of the solar signal (bottom curves

in Figure 4-5), the residual tropical ozone shows a 1-2%

decline in ozone during ~1992-1993, followed by an

increase back to pre-1990 values. Significant long-term

trends in tropical total ozone have not been observed for

the period 1979-2001. H.J. Wang et al. (2002) find weak

negative trends (–1.2 ± 0.8%/decade) in the tropics (20°N20°S) using vertically integrated Stratospheric Aerosol

and Gas Experiment (SAGE) II ozone (15-50 km) for

1985-2000. Although these trends are different from the

near-zero trends derived for column ozone, uncertainty

levels do overlap. 

Figure 4-6 shows deseasonalized ozone values integrated over the 35°-60°S and 35°-60°N latitude belts.

These time series show that ozone levels in the late 1990s

are several percent lower than in the late 1970s in the midlatitudes of both hemispheres, although there is a large

degree of year-to-year variability in both cases. Smoothed

versions of these time series are shown in Figure 4-7.

Time series over NH midlatitudes show an overall

decrease in ozone beginning in the early 1980s, pronounced negative ozone anomalies during ~1992-1996,

and a partial rebound thereafter. The minimum from

~1992-1996 in the smoothed time series (Figure 4-7) is

associated with several individual winters of low ozone

over midlatitudes during that time, namely 1991/1992,

1992/1993, and 1994/1995. The large anomalies in

1991/1992 and 1992/1993 have been associated with the

Mt. Pinatubo volcanic eruption in June 1991 (e.g.,

Brasseur and Granier, 1992; Gleason et al., 1993;

Hofmann et al., 1994).

Total ozone changes in the midlatitude SH also

show a general decline after ~1980 (Figure 4-7), but the

time series show a more consistent and gradual decline

than in the NH. Net ozone levels in the SH midlatitudes

are currently ~6% lower than pre-1980 values. It is especially interesting that the large negative anomalies

observed in the NH extratropics during 1992-1996 are not

particularly evident in the SH. Furthermore, there are relatively large negative anomalies in the SH during ~1985-

1986, which lead to anomalies in the smoothed time series

extending over several years (Figure 4-7), with the result

that the low-frequency midlatitude variability is quite different between the hemispheres. 

4.2.2.2 SEASONAL CHANGES IN TOTAL OZONE

The seasonal changes of total ozone changes over

35°-60°N and 35°-60°S are shown in Figure 4-8, for the

time period 1979-2001. A high degree of year-to-year

variability is observed during the dynamically active

winter and spring periods in each hemisphere: DecemberJanuary-February (DJF) and March-April-May (MAM)

in the NH; June-July-August (JJA) and SeptemberOctober-November (SON) in the SH. In the NH the

largest absolute changes are observed during winter (DJF)

and spring (MAM), with maximum deviations during the

early to middle 1990s of –6 to –12%. During NH summer

(JJA) and fall (SON) the interannual changes show similar time variation (with maximum anomalies during

~1992-1997), but the magnitudes of ozone changes are

less (about –3 to –5%). Total ozone changes in the SH

midlatitudes in Figure 4-8 show a systematic decline

during all seasons, and relatively less seasonal dependence compared with the NH. Although natural variability

is highest during SH winter (JJA) and spring (SON), there

is relatively little difference for long-term changes

between these active seasons and SH summer-fall. Thus

the characteristic seasonal signatures of long-term ozone

loss are distinct between the NH and SH. Seasonally

varying trend calculations are consistent with these time

series results, and results for NH and SH midlatitudes are

shown in Section 4.5.3.

GLOBAL OZONE

4.10
GLOBAL OZONE

4.11

1965 1970 1975 1980 1985 1990 1995 2000

- 4

- 2

0

Ground-based data

TOMS zonal means

SBUV-SBUV/2

Merged satellite data

NIWA assimilated data

1965 1970 1975 1980 1985 1990 1995 2000

- 4

- 2

0 Deviation (%)

Total Ozone Adjusted for Seasonal, QBO, and Solar Effects (smoothed)

60°S-60°N

90°S-90°N Deviation (%)

Figure 4-4. Deseasonalized, area-weighted total ozone deviations, adjusted for solar and QBO effects (as in

Figure 4-3), smoothed in time using four passes of a 13-month running mean to highlight the lowest frequency

interannual variations. Adapted from Fioletov et al. (2002).
GLOBAL OZONE

4.12

- 2

0

2Deviation (%)

0

5 0

100

150

200

250

Ground-based data

TOMS zonal means

SBUV-SBUV/2

Merged satellite data

NIWA assimilated data

1978 1980 1982 1984 1986 1988 1990 1992 1994 1996 1998 2000 2002

- 2

0

2Deviation (%) Solar flux at 10.7 cm

25°S-25°N

Figure 4-5. Upper curve: The smoothed solar flux at 10.7 cm, shown as a proxy for solar cycle variability.

Middle curves: Deseasonalized, area-weighted total ozone deviations from the five datasets for the latitude

range 25°S-25°N. Anomalies were calculated with respect to the ground-based data during 1964-1980.

Lower curves: Deseasonalized data with the effect of the solar cycle removed by statistical regression, and

then smoothed by four passes of a 13-month running mean. Adapted from Fioletov et al. (2002).
GLOBAL OZONE

4.13

1965 1970 1975 1980 1985 1990 1995 2000

-15

-10

-5

0

5 Deviation (%) Ground-based data

TOMS zonal means

SBUV-SBUV/2

Merged satellite data

NIWA assimilated data

1965 1970 1975 1980 1985 1990 1995 2000

-15

-10

-5

0

5 Deviation (%)

Total Ozone Adjusted for Seasonal Effects

35°S-60°S

35°N-60°N

Figure 4-6. Deseasonalized, area-weighted total ozone deviations from five datasets for the latitude bands

35°N-60°N (top) and 35°S-60°S (bottom). Anomalies were calculated with respect to the time average during

1964-1980. Adapted from Fioletov et al. (2002).
GLOBAL OZONE

4.14

1965 1970 1975 1980 1985 1990 1995 2000

- 8

- 6

- 4

- 2

0

2

4

1965 1970 1975 1980 1985 1990 1995 2000

- 8

- 6

- 4

- 2

0

2

4 Deviation (%) Ground-based data

T OMS zonal means

SBUV-SBUV/2

Merged satellite data

N IW A assimilated data

Total Ozone Adjusted for Seasonal Effects and Smoothed

35°S-60°S

35°N-60°N Deviation (%)

Figure 4-7. Deseasonalized, area-weighted total ozone deviations for the midlatitude regions of 35°N-60°N

(top) and 35°S-60°S (bottom) (as in Figure 4-6), but smoothed by four passes of a 13-point running mean.

Adapted from Fioletov et al. (2002).
GLOBAL OZONE

4.15

1980 1985 1990 1995 2000

-12

- 8

- 4

M ar-Apr-May

980 1985 1990 1995 2000

Ground-based data

TOMS zonal means

SBUV-SBUV/2

Merged satellite data

NIWA assimilated data set -12

- 8

- 4 Deviation (%)

D ec-Jan-Feb

1980 1985 1990 1995 2000

-12

- 8

- 4

Sep-Oct-Nov

980 1985 1990 1995 2000

-12

- 8

- 4 D ev iation (%)

Jun-Jul-Aug

1980 1985 1990 1995 2000

-12

-8

-4

4

Mar-Apr-May

980 1985 1990 1995 2000

-12

- 8

- 4

4 D eviation (%)

D ec-Jan-Feb

1980 1985 1990 1995 2000

-12

-8

-4

4

Sep-Oct-Nov

980 1985 1990 1995 2000

-12

- 8

- 4

4 D eviation (%)

Jun-Jul-Aug

35°N-60°N

35°S-60°S

Figure 4-8. Area-weighted total ozone deviations from the 1964-1980 means, calculated for four seasonal

averages, for the latitude bands 35°N-60°N (top) and 35°S-60°S (bottom). Adapted from Fioletov et al. (2002).
GLOBAL OZONE

4.16

4.2.3 Changes in the Vertical Profile of

Ozone

Changes in the vertical profile of ozone are a key

part of the overall fingerprint that helps constrain the

understanding of long-term ozone changes. The SPARC

(1998) and WMO (1999) Assessments demonstrated

that long-term changes in the ozone profile (for the

period 1979-1996) have local maxima in the upper stratosphere (centered near 40 km) and lower stratosphere

(15-20 km), with a relative minimum around 30 km. A

similar result has been discussed by Smyshlyaev and

Geller (2001) based on SAGE II measurements incorporated into a data assimilation model. The analyses

here focus on updated ozone changes in the upper and

lower stratosphere, using both time series and updated

trend analyses (the latter are useful to highlight spatial

patterns of change in sparse datasets, although the

observed time variations are not linear). SAGE I+II

data provide near-global measurements covering both

regions, whereas the Umkehr and ozonesonde data provide independent information primarily over NH midlatitudes. Trend calculations reported here use standard

statistical analyses (SPARC, 1998).

4.2.3.1 UPPER STRATOSPHERE

SAGE I+II

A meridional cross section of the trends derived

from the combined SAGE I (1979-1981) and SAGE II

version 6.1 (v6.1) (1984-2000) data is shown in Figure 4-

9. The overall pattern is very similar to previous SAGE

I+II results that are based on shorter time records (e.g.,

the 1979-1996 results in WMO (1999)). The largest percentage changes are observed in the upper stratosphere

(~35-45 km); maxima are in middle-high latitudes of both

hemispheres, with a magnitude of about –5 to –7%/decade

for the period 1979-2000. The 2s uncertainty levels for

these trends are of order ±1-3%/decade, depending on

location. These trend maxima are slightly less than the

WMO (1999) estimates, which are based on SAGE I+II

data for 1979-1996. The region of significant negative

trends in the upper stratosphere extends throughout the

tropics, which is a change from previous results based on

shorter time records (SPARC, 1998; WMO, 1999). There

are no statistically significant differences between the

upper stratospheric trends in the two hemispheres.

Trends derived from the SAGE II v6.1 data record

alone (for the shorter period 1984-2000) are shown in

SAGE I/II O3 Trends (%/decade) 1979-2000

-60 -40 -20 0 20 40 60 Latitude

20

25

30

35

40

45

50 Altitude (km)

 -9 

 -8 

 -7 

 -7 

 -6 

 -6 

 -6 

 -5 

 -5 

 -5 

 -5 

 -4 

 -4 

 -4 

 -3 

 -3 

 -3 

 -2 

 -2 

 -2 

 -1 

 -1 

 0 

Figure 4-9. Meridional

cross section of ozone

profile trends derived from

the combined SAGE I

(1979-1981) and SAGE II

(1984-2000) datasets.

Trends were calculated in

percent per decade, relative to the overall time

average. Shading indicates that the trends are

statistically insignificant at

the 2s (95%) level.

Updated from H.J. Wang

et al. (2002).
GLOBAL OZONE

4.17

Figure 4-10. The spatial patterns of the trends are similar to the SAGE I+II record (Figure 4-9), but with smaller

negative trends. Significant upper stratospheric trends

for the 1984-2000 period extend into the tropics over 35-

40 km. Upper stratospheric trends in the SH extratropics

are somewhat more negative than those in the NH, but

the differences are not statistically significant. At 40 km,

the 30°-50°S trends are –6.9 ± 2.2%/decade, while those

over 30°-50°N are –6.2 ± 2.2%/decade (2s uncertainty).

Time series of the SAGE I+II record in the upper stratosphere over 40°-50°N are shown in Figure 4-11 (integrated over 38-43 km, corresponding to Umkehr layer

8). The SAGE II data alone show a step-like change

between ~1991 and 1993, with relatively constant ozone

over ~1984-1991 and over ~1993-2000. Similar behavior

is observed from SAGE data in SH midlatitudes (not

shown). Inclusion of the SAGE I data (for 1979-1981)

suggests an overall trend of approximately –5 to –7% per

decade. Note that while the percentage changes in ozone

are relatively high in the upper stratosphere (Figures 4-9

and 4-10), the associated contribution to the total ozone

column change is relatively small (the changes over

1979-2000 in Figure 4-11 are less than 2 Dobson units

(DU), compared with NH midlatitude column ozone

changes of ~10-15 DU for the same period).

Umkehr

Umkehr ozone profile data have recently been reanalyzed using a new retrieval algorithm (Petropavlovskikh

et al., 2001), and these data have been discussed in detail

by Bojkov et al. (2002). Although historic Umkehr data

are available from 13 sites, only a few have continuous

long-term records free from significant discontinuities.

Time series and trend analyses of Umkehr data here are

estimated using the combination of measurements from

Arosa, Switzerland (47°N), Boulder, Colorado, U.S.

(40°N), Haute-Provence, France (44°N), and Belsk,

Poland (52°N). The vertical structure of trends for the

time period 1979-1999 in Figure 4-12 shows a maximum

in the upper stratosphere (layer 8 trends of –5 to

–7%/decade), a minimum near layer 5, and a separate

maximum in the lower stratosphere. The vertical structure and magnitude of the trends are consistent with the

SAGE I+II results in Figure 4-9. Umkehr trends for the

shorter period corresponding to SAGE II (1984-1999) are

smaller than those for 1979-1999 (layer 8 trends of –4 to

–5%/decade), similar to the differences between the

SAGE I+II and SAGE II-only results.

A time series of the deseasonalized Umkehr anomalies in Layer 8 (38-43 km) is included in Figure 4-11 for

SAGE II O3 Trends (%/decade) 1984-2000 

-60 -40 -20 0 20 40 60 Latitude

10

15

20

25

30

35

40

45

50 Altitude (km)

 -7 

 -6 

 -6 

 -5 

 -5 

 -4 

 -4 

 -4 

 -3 

 -3 

 -3 -3 

 -3 

 -2 

 -2 

 -2 -2 -2 

 -2 

 -2 

 -1 

 -1 

 -1 

 -1 

 0 

 0 

 0 

 1 1 

 1 2 2 

 2 

 3 4 5 

Figure 4-10. Meridional

cross section of ozone

profile trends derived

from SAGE II v6.1 data

for the period 1984-2000.

Trends were calculated in

percent per decade, relative to the overall time

average. Shading indicates that the trends are

statistically insignificant

at the 2s (95%) level.

Updated from H.J. Wang

et al. (2002).
GLOBAL OZONE

direct comparison with the SAGE I+II time series. There

is a drop in the Umkehr ozone anomalies between 1982

and 1984 (of order ~ –10% background values), and a similar decrease is found in Nimbus-7 Solar Backscatter

Ultraviolet (SBUV) spectrometer data covering 1979-

1990 (Figure 3.12 of SPARC, 1998). There is a period of

relatively constant Umkehr values over ~1984-1994, and

slight decreases after ~1995. These latter decreases are

not obvious in the SAGE II data, but direct comparisons

of details between these time series are difficult due to the

vastly different space-time sampling.

4.2.3.2 LOWER STRATOSPHERE

SAGE I+II

The SPARC (1998) analyses suggested that the

SAGE I data are only of sufficient quality for trend studies

above 20 km, so trend estimates with the combined SAGE

I+II data cover 20-50 km. Trend estimates for 1979-2000

(Figure 4-9) show significant negative trends over the 20-

to 25-km altitude region in the extratropics of both hemispheres, with magnitudes of –1 to –3%/decade. These

trends are somewhat larger in the NH. The SAGE II data

allow analyses to extend downward in altitude to the

region of the tropopause (~10 km in the extratropics and

~16 km in the tropics). For the SAGE II period of 1984-

2000, significant negative trends of order –2%/decade are

observed in the lower stratosphere over approximately

20-25 km (Figure 4-10). However, significant extratropical trends are not found below ~20 km in either hemisphere for the 1984-2000 record. 

Ozonesondes

Long records of ozonesonde observations (time

series that extend back to at least the early 1970s) are

available for 14 stations covering the latitude region 32°-

75°N (SPARC, 1998). Detailed trend analyses of these

data for the periods 1970-1996 and 1980-1996 are discussed in SPARC (1998), Logan et al. (1999), and WMO

(1999). For both time periods these results reveal significant negative trends that maximize in the lower stratosphere (over pressure levels ~200-50 hPa, or altitudes

~12-20 km). Trends for the period 1980-2000 have been

compared with those for 1980-1996 (based on stations in

the latitude range 36°-59°N), and there are notable differences in the trends with the inclusion of 4 additional years

of data (Figure 4-13). The difference in trends is largest

in the lower stratosphere over ~200-80 hPa (where the

1980-1996 trends are most negative): the mean decrease

in stratospheric ozone for 1980-2000 is 5%/decade at 80

hPa, 2.5%/decade smaller than that for 1980-1996. In

contrast, there is little change in the trend results for pressure levels near and above 50 hPa (altitudes above ~21

km). Comparison of trends for the two periods 1970-1996 Dobson Units

Year

Umkehr

SAGE I+II

Figure 4-11. Top: Time series showing deseasonalized ozone anomalies over latitudes 40°N-50°N,

derived from SAGE I+II data integrated over altitudes 38-43 km. Bottom: Time series derived from

an average of Umkehr measurements at Arosa,

Boulder, Belsk, and Haute-Provence, for data in

Umkehr layer 8 (approximately 38-43 km).

Umkehr ozone trend height (km)

%/decade

50

40

30

20

10 Umkehr layer

Figure 4-12. Vertical profile of ozone trends derived

from Umkehr measurements at Arosa (1979-1999),

Haute-Provence (1984-1999), Boulder (1979-1999),

and Belsk (1979-1999). The dark line indicates the

average trend. The 2s (95%) uncertainty levels are

near ±3%/decade throughout the profile.

4.18
GLOBAL OZONE

and 1970-2000 gives results with similar vertical structure (not shown), but the magnitude of the differences is

much less for the longer record.

The reason for the change in the magnitude of the

trends as a function of altitude can best be understood by

inspection of time series for monthly ozone anomalies.

Figure 4-14 shows deseasonalized ozone anomalies for

four ~3-km layers spanning the lower to middle stratosphere, derived from the average measurements at three

European stations: Hohenpeissenberg, Germany (48°N);

Uccle, Belgium (51°N); and Payerne, Switzerland (47°N).

Time series at the upper levels (63-40 hPa and 40-25 hPa)

show a relatively continuous ozone decline throughout

the record. In contrast, ozone in the lower stratosphere

(the 158-100 hPa and 100-63 hPa levels) shows a decrease

through the early 1990s, with relatively constant levels

thereafter. Similar time series results are found for other

locations. This flattening of the time series after ~1995

results in the reduced trends in the lower stratosphere for

the 1980-2000 period compared with 1980-1996.

The longest continuous ozonesonde record in SH

midlatitudes is from Lauder, New Zealand (45°S), covering the period 1986-2000. Figure 4-15 shows a time

series of these data in 3-km layers, with similar format to

the NH time series in Figure 4-14. While there is substantial interannual variability in these time series (much of it

apparently related to the QBO), there are no significant

stratospheric trends observed in the Lauder data for 1986-

2000. This is somewhat different from the SAGE II trends

in Figure 4-10 (showing negative trends over ~22-27 km

near the latitude of Lauder), but direct comparisons are

difficult in light of the very different space-time sampling.

Note also that the lack of trends at Lauder for 1986-2000

is not inconsistent with the column ozone decreases seen

in Figure 4-6, because much of the column decrease

occurs between the pre- and post-1984 periods.

4.2.4 Vertical Integral of the Profile Trends

vs. Column Trends

The ozonesonde results show that the largest contribution to interannual column ozone changes (over NH

midlatitudes) occurs in the lower stratosphere, over

approximately 160-40 hPa (13-23 km). To quantify that

result, and directly compare the profile-versus-column

ozone changes over NH midlatitudes, Table 4-1 shows the

80-00

80-96

-20 -10 0 10

1000

500

200

100

50

20

10

7 Pressure (hPa)

Trend %/decade -2 -1 0 1

1000

500

200

100

50

20

10

7

0

4

8

12

16

20

24

28

32 Altitude (km)

Trend DU/km/decade

Figure 4-13. Annual mean profile of ozone trends

for eight sonde stations located between 36°N and

59°N, for the time periods 1980-1996 and 1980-

2000. Trend results are shown in percent per

decade (left) and (DU/km/decade) (right). The 2s

error bars are shown.

Table 4-1. Comparison of vertically integrated ozone profile trends (DU/decade) over NH midlatitudes

derived from ozonesondes (an average of eight stations over 36°-59°N), zonal mean SAGE I+II (35°-

60°N), and zonal mean column ozone (35°-60°N), for the period 1980-2000. Results are calculated for

seasonal and annual means. Column ozone trends are calculated from the merged satellite data. The 1s

uncertainty estimates are included in parentheses.

Altitude Ozone Profile Trends (DU/decade)

DJF MAM JJA SON Annual

25-50 km (SAGE I+II) –3.8 (0.9) –2.0 (0.9) –0.9 (0.9) –2.1 (0.9) –1.7 (0.5)

10-25 km (ozonesondes) –8.5 (3.3) –13.2 (2.8) –6.8 (1.8) –4.1 (1.6) –8.2 (1.3)

0-10 km (ozonesondes) +0.9 (0.6) –0.2 (0.9) –1.2 (0.9) –0.2 (0.5) –0.2 (0.4)

Column (merged satellite) –9.0 (3.0) –10.9 (2.8) –5.9 (2.0) –4.0 (1.7) –7.4 (1.2)

4.19
vertical integral of the profile trends derived from

ozonesondes (integrated over 10-25 km), compared with

the zonal mean trends derived from the merged satellite

data (these results are very similar to trends derived from

the other column ozone datasets, as shown in Fioletov et

al. (2002)). Also included are the integrated trends from

SAGE I+II (over 25-50 km), to quantify the contribution

of US ozone losses to the column trends, and tropospheric

ozone trends calculated from ozonesondes integrated over

0-10 km. Trends are calculated from all data for the period

1980-2000 (to directly compare with the ozonesonde

results), and both seasonal and annual mean results are

shown.

The integrated profile and column ozone trends in

Table 4-1 show a number of important features. Ozone

trends in the upper stratosphere (>25 km) derived from

GLOBAL OZONE

4.20 Ozone (DU)

Year

Hohenpeissenberg, Uccle, and Payerne

Smoothed Deseasonalized Monthly Means

Figure 4-14. Time

series of deseasonalized ozone anomalies

for four altitude (pressure) layers, each

approximately 3 km

thick, based on

ozonesonde measurements at three

European stations

(Hohenpeissenberg,

Uccle, and Payerne).

Anomalies were calculated with respect to

1970-1971.
GLOBAL OZONE

4.21

SAGE I+II data are relatively small (of order ~ –2-3

DU/decade), and account for a small fraction (15-30%) of

the observed column losses. Tropospheric ozone trends

do not contribute substantially to the column. There is

reasonable agreement between the integrated ozonesonde

trends (over 10-25 km) and column trends in terms of seasonality (largest losses in NH winter-spring), but less

agreement in terms of magnitude. Partial column trends

for 10-25 km calculated from the ozonesonde data are

consistently larger than the total column trends, for most

seasons and for the annual mean. The agreement with the

total column trends is even worse if the ozone losses above

25 km are included. The reason for the disagreement is

not understood, but is likely related to the sparse sampling

of the ozonesondes (only eight stations are included in

these calculations). Note that the individual ozonesonde

profile integrals are constrained to match co-located

Dobson measurements, and hence the integrated profile

1986 1988 1990 1992 1994 1996 1998 2000

Year

-15

-10

- 5

0

5

1 0

-15

-10

- 5

0

5

1 0

-15

-10

- 5

0

5

1 0 Ozone (DU) -15

-10

- 5

0

5

1 0

158 - 100 hPa

100 - 63 hPa

63 - 40 hPa

40 - 25 hPa

Figure 4-15. Time series of

deseasonalized ozone

anomalies for four altitude

(pressure) layers, each

approximately 3 km thick,

based on ozonesonde

measurements at Lauder,

New Zealand.
trends will necessarily match the Dobson trends for these

eight stations. The fact that the aggregated ground-based

data from ~50 stations agree well with the zonal mean

satellite data (e.g., Figure 4-6) suggests spatial sampling

is at the root of the quantitative imbalances in Table 4-1.

In any case, these results show that the major part of NH

midlatitude column ozone trends is due to loss in the lower

stratosphere (below 25 km), as stated in all previous

Assessments.

4.2.5 Tropospheric Ozone

Trends in tropospheric ozone were discussed in

detail in the previous Assessment (WMO, 1999), and here

we summarize the key points from that report and provide

a brief update. Trends in tropospheric ozone are highly

variable and depend on region and on the time period considered (e.g., Logan et al., 1999; Oltmans et al., 1998;

Tarasick et al., 1995). There were decreases or zero trends

at the Canadian stations for 1970-1996, and decreases of

–2 to –8%/decade for the mid-troposphere for 1980-1996.

The three European stations showed increases for 1970-

1996, but trends were close to zero for two stations for

1980-1996, and positive for one. There were increases in

ozone for the three Japanese stations for 1970-1996, but

trends were either zero or positive for 1980-1996. Time

series of ozone given in Logan et al. (1999) and Oltmans

et al. (1998) show that much of the increase in ozone in

the middle troposphere over Europe occurred before the

mid-1980s.

Trends have been calculated for 1970-2000 and for

1980-2000 and compared with those given by Logan et al.

(1999) for the periods ending in 1996. The addition of 4

years of data makes the tropospheric trends for 1970-2000

less negative for the Canadian stations, and makes the

trends statistically insignificant at Edmonton and Churchill

(these stations have data starting in 1973). The extra 4

years of data make the trends for the European and

Japanese stations slightly less positive. For the period 1980

to 2000, there is no significant trend in mid-tropospheric

ozone for the Canadian stations, the U.S. stations, two

Japanese stations, and one European station, Uccle;

Hohenpeissenberg, Germany, gives a small decrease in

tropospheric ozone, about 4%/decade, and Payerne,

Switzerland, an increase of about 8%/decade. The lack of

trend shown at most stations for 1980-2000 is responsible

for the zero mean tropospheric trend shown in Figure 4-13

for the stations located between 36°N and 59°N.

Updated time series for tropospheric ozone (integrated over 1-10 km) are shown for three stations, in

Europe, Canada, and Japan, in Figure 4-16. At

Hohenpeissenberg, ozone has been lower in the past few

years than the values reached in the mid-1980s and in

1996. All three European stations show continuing ups

and downs in the 1990s, but these are not correlated among

the stations (unlike for the stratosphere, where the

behavior among the three stations is highly correlated).

The Canadian stations show interannual variability in the

1990s, but little long-term changes (e.g., Goose Bay in

Figure 4-16); at most stations, values are lowest in 1992

and 1993, as they are in the stratosphere. Tarasick et al.

(2000) find an increase in tropospheric ozone at all six

Canadian stations, but the increase has occurred only since

1993. The Japanese stations (e.g., Tateno in Figure 4-16)

show relatively little interannual variability in the 1990s.

These results show that tropospheric ozone trends vary

considerably depending on location and the time period

under consideration.

4.2.6 Solar Cycle Effects on Ozone

Global ozone observations display both a long-term

decline and considerable interannual variability on a

decadal time scale, which is approximately in phase with

the 11-year solar cycle (Figures 4-2 and 4-5). In order to

accurately characterize decadal variations in global ozone,

and most directly compare them with model results that

do not include solar forcing, it is important to quantify the

solar cycle effect in the observational record. However,

this is difficult for two reasons. First, the global observational record is relatively short, containing only two or

three solar cycles (somewhat more for ground-based

measurements at a few isolated locations; e.g., Angell

(1989)). Second, a complicating factor for the recent

record is that the large volcanic eruptions of El Chichón

(1982) and Mt. Pinatubo (1991) occurred near the

declining phases of the solar cycle in 1982-1984 and 1992-

1994, and therefore there is the possibility of aliasing the

volcanic and solar effects on ozone (Solomon et al., 1996).

In this section the structure of the solar cycle in the

updated data record is analyzed, including the most recent

solar maximum period during 1999-2001. The availability of data from 1995-2001 provides a new opportunity to distinguish solar from volcanic signals better than

was previously possible, as this is a period of no new volcanic activity together with an additional solar maximum.

Hence, we present an updated analysis that uses the added

years to examine the structure and statistical significance

of the estimated solar signal in ozone, and we discuss comparisons with idealized model simulations.

4.2.6.1 COLUMN OZONE

Changes in solar ultraviolet (UV) spectral irradiance directly modify the production rate of ozone in the

upper stratosphere (e.g., Huang and Brasseur, 1993), and

GLOBAL OZONE

4.22
GLOBAL OZONE

hence it is reasonable to expect a solar cycle variation in

total ozone amount. Analyses of ground-based records

extending over three to six decades indicate the existence

of a decadal time scale variation in column ozone, which

is approximately in phase with the solar cycle (e.g.,

Angell, 1989; Miller et al., 1996; Zerefos et al., 1997).

The global satellite records since 1979 show evidence for

a decadal oscillation of column ozone with large amplitude at low latitudes (Chandra and McPeters, 1994; Hood

et al., 1997; see also Figure 4-5). During at least the last

three solar cycles, the decadal oscillation has been approximately in phase with proxies for solar ultraviolet flux,

and statistical representation is typically based on correlation with the 10.7-cm solar radio flux (F10.7).

The latitudinal structure of the solar cycle variation in column ozone derived from several updated

datasets is shown in Figure 4-17. Here the solar cycle

amplitude is derived from standard regression analyses

(e.g., WMO, 1990), using F10.7 as a proxy for solar variability. These results are derived for several updated data

including (1) the merged satellite data for 1979-2001, (2)

the combined SBUV-SBUV/2 data, (3) the ground-based

data for 1979-2001, and (4) the ground-based data for

1964-2001. Results based on the other satellite datasets

4.23 Ozone (DU)

Year

Ozone Column Anomalies for 1-10 km

Figure 4-16. Time series of

deseasonalized ozone

anomalies for tropospheric

column ozone, integrated

between 1 and 10 km, for

Hohenpeissenberg, Germany; Goose Bay, Canada;

and Tateno, Japan. The

dashed lines in the Goose

Bay time series denote

measurements by Brewer

Mast ozonesondes, which

were replaced by electrochemical concentration cell

(ECC) measurements (solid

lines).
GLOBAL OZONE

4.24

for 1979-2000 (which rely heavily on Total Ozone

Mapping Spectrometer (TOMS) data) are similar to those

from the merged satellite data and are not shown here.

The derived solar cycle amplitude in Figure 4-17 from the

satellite datasets shows a relatively flat latitudinal structure, with amplitudes in the tropics of approximately 5

DU/(100 units of F10.7) for the merged satellite data and

3.5 DU/(100 units of F10.7) for the SBUV-SBUV/2 data;

the larger amplitude in the merged satellite data may be

an artifact of the Nimbus-7 TOMS data, as discussed in

Appendix 4A. Note that F10.7 varies from ~75 to ~200

units over a solar cycle (Figure 4-5), and therefore these

derived amplitudes should be multiplied by ~1.25 to

obtain the corresponding column ozone change over a

solar cycle (tropical column ozone variations work out to

approximately 2-3% between solar maximum and solar

minimum). The statistical uncertainty of these fits is relatively large, with 2s uncertainty levels of order ±1-2

DU/(100 units of F10.7) or larger, and the regression fits

are most statistically significant in the tropics (~30°N30°S).

The solar cycle amplitudes derived from the

ground-based data for both the 1964-2001 and 1979-2001

time periods are similar to each other, and have significantly more latitudinal structure than seen in the satellite

data results. Results over the NH (where the ground-based

data are most dense) show values roughly consistent with

the range of satellite-derived values. However, in the less

well sampled SH, the ground-based solar signal is smaller

than that derived from the satellite data, ranging from 2 to

3 DU/(100 units F10.7). The similarity of results from

the two time periods (1964-2001 and 1979-2001) shows

that an increase of sampling from two solar cycles to three

does not strongly influence these empirical estimates.

Solar cycle variations in ozone have been studied

in a number of idealized modeling studies (e.g., Huang

and Brasseur, 1993; Fleming et al., 1995; Lee and Smith,

2002). While this is (conceptually) a straightforward calculation, given knowledge of the variations in solar spectral irradiance over a solar cycle (e.g., Lean et al., 1997),

there are possible dynamical feedbacks in the stratospheric

climate system that can complicate the full solar influence. In addition, because there is likely some confusion

between direct solar forced and volcanic effects on ozone

for recent decades (as demonstrated in the idealized model

results in SPARC (1998)), it is useful to perform model

simulations with and without realistic aerosol variations

to study this influence. 

The simulated solar cycle in column ozone shows

similarity between different model results, with details

depending on the specific model. Fleming et al. (1995)

and Lee and Smith (2002) derive a solar signal in column

ozone with broad latitudinal structure, and a solar

maximum-to-solar minimum amplitude of 1.5-2% (corresponding to 3-5 DU / (100 units of F10.7)), which is in

reasonable agreement with observations in Figure 4-17.

Lee and Smith (2002) include realistic aerosol effects for

1979-2000, which slightly increases the apparent solar

signal in their model, but primarily outside of the tropics.

4.2.6.2 VERTICAL PROFILE OF THE SOLAR SIGNAL

The dependence on altitude of the ozone solar cycle

has been investigated using SAGE I+II data by H.J. Wang

-60 -30 0 30 60

Latitude

- 2

0

2

4

6

8A mplitude (DU/100 Units F10.7 Flux)

Ground-based data, 1979-2001

Ground-based data, 1964-2001

SBUV-SBUV/2

Merged satellite data

Figure 4-17. Latitudinal structure of the solar cycle regression

coefficient calculated from the

merged satellite data, SBUV+-

SBUV/2 data, and ground-based

data for the period 1979-2001,

and also from the longer record

of ground-based data over 1964-

2001. Error bars show 2s

regression uncertainty estimates

for the SBUV+SBUV/2 results,

and uncertainties for the other

data are similar (not shown).
GLOBAL OZONE

4.25

et al. (1996) and SPARC (1998). The results show

coherent, in-phase ozone variations in the upper stratosphere (~35-45 km), with local variations of ~2-4%/(100

units F10.7), and these values are not substantially

changed by the update of the SAGE II data through 2000

(not shown). Similar overall results are obtained from the

shorter record of SBUV+SBUV/2 data for 1979-1994

(Chandra and McPeters, 1994; SPARC, 1998), and from

analyses of Umkehr data over NH midlatitudes (Miller et

al., 1996). The results from both SAGE I+II and

SBUV+SBUV/2 measurements show an upper stratospheric solar signal with a minimum in the tropics and

maxima in midlatitudes of both hemispheres, which is different from the broad latitudinal structure derived in idealized model simulations (e.g., Huang and Brasseur,

1993). Lee and Smith (2002) have suggested that this pattern arises in the short satellite observational record

because of aliasing with volcanic and possibly QBO

effects. The SAGE I+II and SBUV+SBUV/2 analyses

also suggest a solar signal in the tropical lower stratosphere, which is not simulated in idealized models;

SPARC (1998, Section 3.4.2) suggests that signal may

arise from volcanic aliasing effects.

4.2.6.3 SUMMARY

Both global observations and idealized model simulations show column ozone variations that are approximately in phase with the 11-year solar cycle. The calculations of Lee and Smith (2002) suggest that volcanic

aerosol effects alias onto the solar cycle in column ozone

for the 1979-2000 period, but the effect is not excessive

in the tropics (<30%), where solar effects are most significant in observations. The amplitude of the modeled

response in the tropics is broadly within the range of

observations in Figure 4-17, i.e., ~3-5 DU/(100 units

F10.7), although there is considerable uncertainty in

details of the observed “signal.” In spite of these uncertainties, the decadal variability in column ozone observed

in the tropics (and in the global mean) is reasonably consistent with modeled solar variability, and model results

suggest that the effect can be approximately isolated

during 1979-2000 based on regression analyses. In contrast, there may be large uncertainties in isolation of the

solar signal in the extratropics.

4.3 CURRENT UNDERSTANDING OF PAST

CHANGES IN STRATOSPHERIC AEROSOL,

WATER VAPOR, AND NO2

Important components of the stratospheric climate

system that directly influence ozone (and that are not discussed explicitly in Chapter 1) include stratospheric

aerosol, water vapor, and nitrogen dioxide (NO2). This

section discusses the updated observations and current

understanding of long-term changes in these constituents.

4.3.1 Stratospheric Aerosol Variability

Since the late 1970s, stratospheric aerosols have

been observed by a variety of ground-based, in situ, and

space-based instruments. In this two-decade time period,

volcanic eruptions have strongly influenced stratospheric

aerosol loading. The eruptions include those of El

Chichón (Mexico, 1982) and Mt. Pinatubo (Philippines,

1991), the latter probably having the largest stratospheric

impact of any volcano since Krakatau in 1883. On the

other hand, stratospheric aerosol loading reached very low

levels in the late 1970s, the late 1980s, and from the late

1990s to the present. On this time scale, stratospheric

aerosol is perhaps the most variable component of the

stratospheric composition. This variability is important

because of the role aerosol plays in stratospheric ozone

change through heterogeneous chemical processes as well

as through aerosol-induced changes in stratospheric circulation.

Figure 4-18 shows a 25-year record of stratospheric

aerosol variability based on integrated backscatter from

the Fraunhofer Institute for Atmospheric Environmental

Research (IFU) lidar (Jäger et al., 1995), together with 17

years of 1020-nm aerosol optical depth measurements by

SAGE II. The dominant features are the 1982 eruption of

El Chichón and the 1991 eruption of Mt. Pinatubo, which

raised the integrated backscatter and aerosol optical depth

values to more than 100 times their minimum values

observed during this period. The effects of several smaller

volcanic eruptions (indicated in the figure) can also be

seen. The volcanic aerosols from the large tropical eruptions of El Chichón and Mt. Pinatubo were mixed relatively rapidly throughout the stratosphere, and time series

over most of the globe are similar to those in Figure 4-18.

With no significant eruptions after 1991, it is clear from

Figure 4-18 that the relaxation to a nonvolcanic aerosol

level lasted at least until 1999, i.e., more than 8 years.

Therefore, with the possible exception of the late 1970s

and the last few years, the entire record can be characterized as continual recovery from episodic volcanic injections of aerosol and their gaseous precursors (Bluth et al.,

1992). This has made it extremely difficult to infer any

long-term trends in aerosols associated directly with

human activities or associated with other climate change

phenomena. Longer-term measurements from individual

stations (such as for the IFU lidar in Figure 4-18) show

that current stratospheric integrated backscatter levels are
about the same or even less than those observed in the late

1970s, the period commonly considered the nominal stratospheric background period. Although various mechanisms have been suggested (e.g., Hofmann, 1990, 1991),

there is currently no compelling evidence of a trend in the

nonvolcanic aerosol loading, at least as it is measured by

bulk parameters such as extinction or backscatter

(Thomason et al., 1997).

The midlatitude records shown in Figures 4-18

reveal a clear annual cycle in addition to the long-term

recovery from volcanic eruptions. In the tropics, the

record mainly reflects monotonic decreases between volcanic events. However, during the recent clean period, an

annual cycle in the column optical depth can be observed.

The Climate Monitoring and Diagnostics Laboratory

(CMDL) lidar record at Mauna Loa, Hawaii, shows an

annual cycle during the current clean period that has been

attributed to a transport of new material introduced into

the tropics (Barnes and Hofmann, 2001).

Unlike instruments that measure ozone and most

other stratospheric species, the instruments that quantify

stratospheric aerosol often measure fundamentally different aerosol properties. Rather than measuring number

density of a molecule, satellite-based instruments and

aerosol lidar measure optical properties that arise from an

underlying size distribution and composition that can only

be indirectly inferred. As a result, assessing the quality of

remotely sensed data and consistency between instruments can be challenging. This problem is illustrated in

comparing the SAGE II 1020-nm stratospheric aerosol

optical depth with the IFU integrated backscatter in Figure

4-18 (with the optical depth divided by 40 for comparison). In most time periods the agreement between the

two records is very good, because both capture the longterm recovery and the annual cycle. However, in the late

1980s the lidar value is nearly 50% larger than the adjusted

SAGE II values. This could reflect the conversion

between optical depth and backscatter (an extinction-tobackscatter ratio of 40), though it does not occur in a similar period in the 1990s, or it could reflect limitations in

one or both instruments. However, the nature of the measurements makes an objective assessment difficult. 

This problem has a direct impact on the assessment

of the effect of aerosol on heterogeneous chemical

processes and radiative/dynamical processes, and

ultimately on ozone. The rates of most heterogeneous

reactions are directly related to the aerosol surface area

density, whereas extinction (i.e., total scattering plus

absorption) is roughly proportional to aerosol volume density. Many models use a climatology of aerosol surface

area and volume densities based on satellite observations

as either direct input (e.g., Section 4.5.3) or as a verification of an internal simulation. Two long-term, global climatologies are derived from SAGE II and the Halogen

Occultation Experiment (HALOE), which both use solar

occultation measurements. SAGE II measures aerosol

extinction in the visible and near-infrared, where extinction is dominated by scattering, whereas HALOE measures extinction in the infrared where extinction is dominated by aerosol absorption. Surface area and volume

GLOBAL OZONE

4.26 Integrated Backscatter (sr-1) 10-2

10-5

10-4

10-3 La Soufiere Sierra Negra St. Helens Alaid Nyamuragira El Chichon Ruiz Nyamuragira Redoubt Pinatubo

Kliuchevskoi Shishaldin

1979

Year

76 78 80 82 84 86 88 90 92 94 96 98 00

Figure 4-18. Multiyear time series of stratospheric aerosols measured by lidar (694.3 nm) at Garmisch (47.5°N,

11.1°E) in Southern Germany (red curve) and zonally averaged SAGE II stratospheric aerosol optical depth

(1020 nm) in the latitude band 40°N-50°N (black curve). Vertical arrows show major volcanic eruptions. Lidar

data are given as particle backscatter integrated from 1 km above the tropopause to the top of the aerosol layer.

The curve referring to SAGE II data was calculated as optical depth divided by 40. For reference, the 1979

level is shown as a dashed line. Data from Garmisch provided courtesy of H. Jäger (IFU, Germany).
GLOBAL OZONE

4.27

density data from these instruments agree reasonably well

with data derived from the in situ optical particle counter

(OPC) (Thomason et al., 1997; Hervig et al., 1997; Hervig

and Deshler, 1998; Hervig and Deshler, 2002), with differences generally within 30 to 50%. An example of this

reasonable agreement is shown in Figure 4-19. On the

other hand, comparisons with measurements of aerosol

surface area density by the in situ Focused Cavity Aerosol

Spectrometer of Denver University are systematically

larger than SAGE II estimates by as much as 100%. These

differences may reflect the surface area of particles with

radii smaller than 100 nm, which are effectively invisible

to SAGE II because of their poor scattering efficiency in

the visible and near-infrared (J.C. Wilson, University of

Denver, personal communication). These differences

between inferred surface area densities are large enough

to introduce significant uncertainties into modeling of

heterogeneous chemistry.

4.3.2 Stratospheric Water Vapor Trends

The recent Stratospheric Processes and their Role

in Climate (SPARC) assessment of upper tropospheric

and stratospheric water vapor (SPARC, 2000) has provided an extensive review of data sources and quality for

stratospheric water vapor, together with detailed analyses

of observed seasonal and interannual variability. The

longest continuous reliable dataset is at a single location

(Boulder, Colorado, U.S.), based on balloonborne frostpoint hygrometer measurements, and dates back to 1981.

Over the period 1981-2000, a statistically significant positive trend of approximately 1%/year is observed at all

levels between about 15 and 28 km (SPARC, 2000;

Oltmans et al., 2000) (Figure 4-20). Increasing water

vapor trends of ~2%/year are also found for 19-28 km in

balloon measurements made from Washington, D.C.,

U.S., during 1964-1976 (Oltmans et al., 2000). However,

although a linear trend can be fitted to these data, there is

a high degree of variability in the infrequent sampling at

individual locations, and the increases in either record are

neither continuous nor steady. Long-term increases in

stratospheric water vapor are also inferred from a number

of other ground-based, balloon, aircraft, and satellite

datasets spanning ~1980-2000 (Rosenlof et al., 2001),

although the sampling uncertainty is high in many cases.

Hurst et al. (1999) did not find a significant trend in the

quantity [H2O+2CH4] in the lower stratosphere from Earth

Resources (ER)-2 aircraft measurements between 1993

and 1997, but this result is based on only eight infrequent

time samples.

0 300 600 900 1200

5

10

15

20

25 Wyoming OPC 

SAGE II Version 6.0

Days after Pinatubo eruption Surface area, μm2 cm-3 Figure 4-19. Comparison of surface area density

inferred from the University of Wyoming optical particle counter (OPC) in Laramie, Wyoming, and that

from SAGE II (zonally averaged between 35°N and

45°N). Figure updated from Thomason et al. (1997).

Figure 4-20. Vertical profile of trends in water vapor

over Boulder, Colorado, U.S., for the period 1981-

2000, expressed in percentage of the long-term

means. The solid line is the trend computed at

250-m intervals, and the shaded region is the 95%

confidence region. From Oltmans et al. (2000).
GLOBAL OZONE

4.28

Global stratospheric water vapor measurements

have been made by the HALOE satellite instrument for

more than a decade (Russell et al., 1993). These data show

long-term increases over almost the entire stratosphere

above 25 km (Nedoluha et al., 1998a; Evans et al., 1998;

Randel et al., 1999; C.A. Smith et al., 2000). For the

period 1991-2001 the HALOE-derived water vapor trends

have a magnitude of ~25 to 45 parts per billion by volume

(ppbv) per year, which corresponds to local increases of

~0.6 to 0.8%/year, slightly smaller than the 1981-2000

Boulder balloon measurements. However, the HALOE

data do not show significant trends in the lower stratosphere (between 20 and 25 km), which is a curious result

given the strong coupling of stratospheric water vapor to

the tropical tropopause region. Time series of the HALOE

data furthermore show relatively large increases during

the period 1991-1996, and a flattening after 1996 (Randel

et al., 1999).

Causes of decadal-scale increases in stratospheric

water vapor are not well understood at present. A substantial fraction of the stratospheric water vapor budget

originates from the oxidation of methane, which accounts

for the increase of climatological water vapor with altitude for 20-50 km. However, the observed ~1%/year

water vapor increase observed in balloon and satellite

observations is more than twice the magnitude that would

be expected from observed trends in tropospheric methane

alone. A simple explanation for the stratospheric water

vapor increase would be an increase in average saturation

mixing ratio at the tropical tropopause. However, observations of tropical tropopause temperature do not show a

corresponding increase, but rather suggest a small

decrease during the period 1979-1998 (Simmons et al.,

1999; Randel et al., 2000; Seidel et al., 2001; Zhou et al.,

2001a). The inability to explain past changes diminishes

our confidence in the projection of future increases in

stratospheric water vapor. 

4.3.3 Trends in Stratospheric NO2

Nitrogen dioxide (NO2) plays a key role in stratospheric ozone chemistry, not only through the NOx (NO +

NO2) catalytic destruction cycle, but also by interacting

with the ClOx (chlorine radicals) and HOx (odd hydrogen)

families, leading to the formation of important stratospheric reservoirs such as chlorine nitrate (ClONO2) and

nitric acid (HNO3). Long-term monitoring of all these

nitrogen compounds is among the priority tasks of the

Network for the Detection of Stratospheric Change

(NDSC; see http://www.ndsc.ws).

The longest record of column-integrated stratospheric NO2 is available from measurements made since

1981 at the primary NDSC station of Lauder, New Zealand

(45°S, 170°E, 370 m above sea level (asl)). The data

(Liley et al., 2000) have been derived from UV-Visible

scanning spectrometers that measure sunlight scattered

from the zenith sky during sunrise and sunset. The time

series at 90° solar zenith angle is continuous from 1981 to

the present except for a break of 6 months in 1986.

Measurement intercomparisons (Hofmann et al., 1995;

Roscoe et al., 1999) have shown that the instruments and

analysis procedure meet the standards of the NDSC. Time

series of deseasonalized NO2 anomalies in the sunrise and

sunset Lauder measurements are shown in the lower panel

of Figure 4-21 for the period 1981-2001. Large episodic

decreases of NO2 are observed following the volcanic

eruptions of El Chichón (1982) and Mt. Pinatubo (1991),

persisting for 2-3 years in each case. Linear-trend statistical fits to these data (Liley et al., 2000) show an upward

trend of 5 ± 1%/decade over the period 1981-1999; trends

extended through 2001 are somewhat larger, i.e., 6 ±

1%/decade.

Midday NO2 column abundances have also been

derived from the analysis of infrared solar spectra

recorded since 1985 at the primary NDSC site of

Jungfraujoch, Switzerland (46.5°N, 8.0°E, 3580 m asl).

These measurements were made with Fourier transform

instruments on clear-sky days during observation campaigns that became more regular and frequent after the

NDSC was formalized in 1989. The deseasonalized

anomalies from the Jungfraujoch measurements (Figure

4-21) show an overall increase with time, together with a

transient NO2 decrease associated with the Mt. Pinatubo

volcanic eruption and its subsequent recovery (De

Mazière et al., 1998), similar to that seen at Lauder.

Trends calculated from the Jungfrauojoch data give a

linear rate of increase of 6 ± 2% per decade for the period

1985-2001 (an update of Mahieu et al., 2000). This trend

is in good agreement with the Lauder results, indicating

that the stratospheric burden of NO2 has been evolving

similarly in midlatitudes of both hemispheres since the

mid-1980s.

The observed trends in stratospheric NO2 above

Lauder and the Jungfraujoch are about twice that of the

total reactive nitrogen (NOy) source gas, nitrous oxide

(N2O), for which surface emissions are known to be

increasing the tropospheric concentration by about 3% per

decade (WMO, 1999; also Chapter 1 of this Assessment).

Model studies (Fish et al., 2000; McLinden et al., 2001)

show that much of this difference is largely a result of

ozone depletion. The model predictions of McLinden et

al. (2001) of a +4.3%/decade trend in NO2 consist of

+2.4%/decade due to the N2O increase, +2.5%/decade due

to ozone depletion, and –0.6%/decade from the halogen
GLOBAL OZONE

4.29

trend’s impact on odd-nitrogen partitioning. Changes in

the vertical distribution also contributed +0.4%/decade to

the trend in slant-column density. However, this study

ignored the effect of H2O increases, which would decrease

the modeled trend by 0.6%/decade (Fish et al., 2000).

Overall the observed N2O increases and decreases in ozone

appear sufficient to explain (within uncertainties) an NO2

trend of 5 ± 1%/ decade.

4.4 CURRENT UNDERSTANDING OF PAST

CHANGES IN STRATOSPHERIC

TEMPERATURE

An evaluation of stratospheric temperature trends

is an integral part of understanding ozone trends. Ozone

is a key radiatively active gas throughout the stratosphere,

and it is important to assess the degree of consistency

between observed ozone trends and temperature trends.

Furthermore, in the lower (upper) stratosphere halogenrelated ozone depletion is generally accelerated (decelerated) by lower temperatures, and therefore quantification

of temperature changes is crucial for understanding past

and predicting future ozone changes.

4.4.1 Updated Observations of

Stratospheric Temperature

An extensive evaluation of stratospheric temperature trends was conducted by the SPARC Stratospheric

Temperature Trends Assessment, with results included in

WMO (1999), along with some updates in Ramaswamy et

al. (2001). We briefly report updated satellite-derived time

series, based on Microwave Sounding Unit (MSU) and

Stratospheric Sounding Unit (SSU) measurements; the

latter are described in detail in WMO (1999). MSU

channel 4 provides a measure of the weighted mean temperature for altitudes of 13-22 km, while several different

channels of SSU provide temperatures for layers spanning

the lower to upper stratosphere, with vertical resolution of

about 10-15 km. We also include brief comparisons with

temperatures from the National Centers for Environmental

Prediction/National Center for Atmospheric Research

(NCEP/NCAR) reanalyses (Kalnay et al., 1996), which

incorporate both satellite and radiosonde measurements in

global analyses. The earlier assessments have discussed

other sources of temperature trend data and the reasonable

agreement between these and the satellite data. Percent

Year

Figure 4-21. Time series of

deseasonalized anomalies

in slant column measurements of nitrogen dioxide

(NO2) measured at Jungfraujoch, Switzerland (46°N),

during 1985-2001, and at

Lauder, New Zealand (45°S),

for 1981-2001. Lauder results are shown for measurements made at sunrise

(dashed lines) and sunset

(solid lines).
GLOBAL OZONE

4.30

Time series of global mean temperature anomalies

at levels spanning the lower to upper stratosphere are

shown in Figure 4-22, based on deseasonalized monthly

mean MSU (1979-2001) and SSU (1979-1998) datasets.

Globally averaged temperatures in the lower and middle

stratosphere (up to about 30 km) show a strong imprint of

volcanic eruptions, with warming for 1-2 years following

El Chichón (1982) and Mt. Pinatubo (1991) eruptions.

Globally averaged temperatures near 17 km show an overall

cooling from 1979 to 2001, but the anomalies are relatively

constant for the period between the large eruptions; the time

series suggest a “step-like” global-scale cooling of approximately 0.5 K following the eruptions. Also included for

the lowest two levels in Figure 4-22 are globally averaged

anomaly time series derived from the NCEP/NCAR

reanalyses for 1979-2000 (100-hPa data plotted at 17 km,

and 50-hPa data at 22 km). There is good overall agreement between the thick-layer satellite measurements and

the reanalyses time series in the lower stratosphere.

The spatial structure of the annual mean stratospheric temperature trends for 1979-1998 is shown in

Figure 4-23. Trends were calculated using a standard

regression analysis, including proxies for the QBO and

solar cycle variations, and the two years following the El

Chichón and Mt. Pinatubo volcanic eruptions were

omitted from the time series to remove any volcanic influences (this has a small effect on the calculated trends, as

shown in WMO (1999)). The annual mean trends show a

high degree of interhemispheric symmetry. In the lower

stratosphere there is statistically significant cooling in

the extratropics (20°-60°N and S) of about 0.4 to 0.8

K/decade, which is relatively constant throughout the year

(not shown). On the basis of satellite data, the equatorial

lower stratosphere exhibits weak cooling (less than 0.4

K/decade), which is not statistically significant. There is

a relative minimum in cooling in the middle stratosphere

(around 30-35 km). The cooling increases monotonically

with height from 35 to 50 km, with maximum values of

about 2.5 K/decade near the stratopause (50 km).

4.4.2 Causes of Past Temperature

Changes

The previous Assessment (WMO, 1999) concluded

that ozone depletion was the dominant cause of globalmean temperature trends in the lower stratosphere. Ozone

Figure 4-22. Time series of global mean

temperature anomalies derived from

satellite radiance measurements. The

upper four curves (solid lines) are from

the Stratospheric Sounding Unit (SSU)

channels 47X, 27, 26X, and 15X (top to

bottom) and represent mean temperatures for 10- to 15-km layers centered

near 50, 42, 30, and 22 km, respectively.

The bottom time series (solid line) is a

global mean from the Microwave

Sounding Unit (MSU) channel 4, which

has a peak contribution near 17 km. The

dashed lines for the two lower levels

show global temperature anomalies

derived from the NCEP/NCAR reanalyses, for pressure levels 50 hPa (22

km) and 100 hPa (17 km). The vertical

dashed lines indicate the volcanic eruptions of El Chichón (1982) and Mt.

Pinatubo (1991).
GLOBAL OZONE

4.31

depletion and increases in the more well-mixed greenhouse gases (WMGGs) both contributed significantly in

the middle and upper stratosphere. It also noted that there

was little evidence that changes in tropospheric climate

and sea surface temperatures have played a significant

role in global-mean stratospheric temperature trends.

There have been significant developments in this

area since WMO (1999). First, at the time of WMO (1999),

most model studies used rather idealized vertical distributions of ozone change; more refined datasets on the vertical trends in ozone are now in widespread use. Second,

several groups have reported simulated temperature trends

using coupled chemistry-climate models. Third, much

attention has been paid to the potential role of stratospheric

water vapor changes on temperature trends. 

The analysis here is centered on an intercomparison of available model results that have appeared in the

literature or that have used established models. The

models range from simple fixed dynamical heating (FDH)

models, to two-dimensional (2-D; latitude-height) models,

to general circulation models (GCMs) that use imposed

ozone trends. The 2-D models and GCMs that incorporate coupled chemistry are discussed in Shine et al. (2002).

Table 4-2 lists the contributors to this intercomparison,

together with some model details. More details of these

intercomparisons were reported by Shine et al. (2002).

The emphasis of this section is on global-scale temperature changes. Chapter 3 considers polar temperature

trends. All plots here include trends from the MSU and

SSU data discussed above (Section 4.4.1). They are

included on all plots for individual temperature trend

mechanisms for reference, but the summary discussion

will bring together the information from these individual

mechanisms into a coherent picture. The abbreviation

WMGG will be used to indicate the longer-lived and more

well-mixed greenhouse gases: carbon dioxide (CO2),

methane (CH4), nitrous oxide (N2O), and the halocarbons.

For the GCM simulations, the 2s unforced variability of

the global-mean temperature is less than 0.25 K.

4.4.2.1 SIMULATED TEMPERATURE TRENDS DUE TO

OBSERVED OZONE TRENDS

Improved analyses of vertical profiles of ozone

change over the period from about 1979 to 1997 (e.g.,

Randel and Wu, 1999; Langematz, 2000) have now been

used in models to re-examine the effect of ozone loss on

temperature trends (Langematz, 2000; Ramaswamy and

Temperature Trend (K/decade) 1979-1998 Height (km)

Latitude Pressure (mb)

60S 30S 0 30N 60N

50

40

30

20

 0

3

10

30

100

Figure 4-23. Meridional cross sections

of trends (K/decade) in

stratospheric temperature during January

1979-May 1998, derived from combined

SSU and MSU satellite

datasets. Shading indicates that the trends

are significant at the

2s level.
GLOBAL OZONE

4.32

Schwarzkopf, 2002; Rosier and Shine, 2000; Smith, 2001;

and Forster et al., 2001, who used three different radiation schemes). Figure 4-24 shows the vertical profile of

the global and annual mean temperature trends as a function of pressure when only ozone changes are imposed.

All models have the same qualitative behavior, with a

peak cooling near the stratopause (~1 hPa) exceeding

1 K/decade, a minimum near 10 hPa, and a secondary

cooling peak near 80 hPa of about 0.4 K/decade. However, there is a significant amount of disagreement in both

the exact size of the cooling and the altitude of the main

features, most particularly near 1 hPa. The dominant

reason for these differences seems likely to be the particular radiation codes being used (Forster et al., 2001) and

the assumed background climatology of ozone

(Ramaswamy and Schwarzkopf, 2002).

The latitudinal structure of annual and zonal mean

temperature trends at 50 hPa from observations and

models with observed ozone trends is shown in Figure 4-

25. It is important to appreciate that whereas dynamical

processes exert little direct influence on global-mean temperature at any given level (in the stratosphere), this is

decidedly not true of temperatures over particular latitude

bands. The general pattern in the GCMs is of a cooling of

Table 4-2. Contributors to the temperature trend intercomparison. “Timeslice” refers to the calculation of

temperature trends by doing steady-state integrations of the model for conditions representative of 2 or more

years and calculating the trend from the difference between these simulations. Transient models perform continuous integrations through the given period. The column labeled C refers to whether CO2 changes are

included, O refers to whether other well-mixed greenhouse gases are included, and H refers to whether water

vapor trends are imposed. Note that in the absence of imposed water vapor trends, most models will have

some water vapor change in response to other perturbations.

Institute Contributor Model Ozone Trend C O H Experiment Reference

Univ. of Cambridge, P. Braesicke GCM Coupled but no N N N Transient, Braesicke et al.

U.K. changes in obs. SSTs (2002)

trace gases

Freie Universität, U. Langematz GCM Observed trend Y N N Timeslice Langematz (2002)

Berlin, Germany

DLR, Germany V. Grewe, GCM Coupled Y N N Timeslice Schnadt et al. (2002)

C. Schnadt

Imperial College, C.A. Smith, 2-D Observed trend N N Y Timeslice C.A. Smith et al.

U.K. J.D. Haigh (2001)

Imperial College C.A. Smith, GCM – N N Y Timeslice Model is Rosier

U.K. J.D. Haigh and Shine (2000)

NOAA Aeronomy R.W. Portmann, 2-D Coupled Y N Y Timeslice Dvortsov and

Laboratory, U.S. S. Solomon Solomon (2001)

NOAA GFDL, U.S. V. Ramaswamy, GCM Observed trend Y Y Y Timeslice Ramaswamy and

M.D. Schwarz- Schwarzkopf 

kopf (2002)

NASA GISS, U.S. D. Shindell GCM Coupled Y Y Y Transient Shindell (2001)

Met Office (UKMO), J. Austin, GCM Coupled Y Y N Transient Austin (2002)

U.K. N. Butchart

Met Office (UKMO), N. Butchart GCM – Y Y N Transient Butchart et al. (2000)

U.K.

Natl. Inst. of Env. T. Nagashima GCM Coupled Y Y N Transient Nagashima et al.

Studies, Japan (2002)

Univ. of Reading, M. Bourqui, GCM Observed trend Y Y N Timeslice Rosier and Shine

U.K. K. Shine (2000)

Univ. of Reading, P. Forster FDH Observed trend N N Y Timeslice Forster et al. (2001)

U.K.

DLR, Deutsches Zentrum für Luft- und Raumfahrt; NOAA, National Oceanic and Atmospheric Administration; GFDL, Geophysical Fluid Dynamics

Laboratory; NASA, National Aeronautics and Space Administration; GISS, Goddard Institute for Space Studies. 
GLOBAL OZONE

4.33

a few tenths of a Kelvin per decade, which is substantially

less than the observed cooling in midlatitudes. At 100

hPa (not shown) the GCMs are in better agreement with

observations: the models exhibit midlatitude cooling

trends of around 0.3 K/decade, which overlaps the uncertainty levels of the (somewhat larger) observed trends. As

a note, there are considerably larger latitudinal gradients

in the temperature trends near Antarctica in the FDH

models compared with the GCMs in Figure 4-25 (with the

exception of the GFDL GCM); as discussed in Rosier and

Shine (2000), changes in the GCM model circulation tend

to ameliorate the radiatively driven high-latitude cooling.

A further contributor to temperature trends in the

lower stratosphere comes from changes in tropospheric

ozone concentrations (WMO, 1999); these change the

amount of upwelling radiation that can be absorbed by

stratospheric ozone. Sexton et al. (2002), using tropospheric ozone increases derived from a chemical transport model (CTM), found a global mean cooling of about

0.05 K/decade at 50 hPa in recent decades. Although

this cooling is smaller than that due to stratospheric

ozone loss, it has likely been sustained over many

decades; Sexton et al. (2002) estimated a 50-hPa cooling

of about 0.5 K from tropospheric ozone changes over

the 20th century.

4.4.2.2 SIMULATED TEMPERATURE TRENDS DUE TO

INCREASED WMGG CONCENTRATIONS

There is qualitative agreement between the modeled cooling resulting from increases in WMGG concentrations. It generally increases monotonically with height

from near zero at 100 hPa to about 0.8 K/decade at the

–3 –2.5 –2 –1.5 –1 –0.5 0

Temperature Trend K/decade

0.01

0.10

1.00

10.00

100.00 Pressure (hPa) Berlin GCM

Reading IGCM

Imp. Coll. (2-D)

Reading FDH (NBM)

Reading FDH (ECHAM)

Reading FDH (Zhong)

GFDL GCM

SSU/MSU Observed Trends (79-97)

Figure 4-24. Global and annual mean temperature trends for models using imposed height-resolved ozone

trends. The observations from MSU and SSU are also shown. The 2s error bars in the observations are

included; the vertical bars are intended to give the approximate altitude range sensed by the particular satellite channel. NBM, ECHAM, and Zhong refer to different radiative transfer codes used in the Forster et al.

(2001) fixed dynamical heating model. NBM is a moderately high-resolution narrow band code; ECHAM and

Zhong are general circulation model codes.
GLOBAL OZONE

4.34

stratopause (not shown, but see Figure 5-24 of WMO

(1999)). Between 100 and 10 hPa, there is general agreement between most models, but there is considerable

divergence at about 1 hPa, with almost a factor of 4 difference between the model with the most cooling (UKMO,

the United Kingdom Meteorological Office model) and

that with the least (GISS, the Goddard Institute for Space

Studies model). Given that the greenhouse gas changes

are fairly linear and well constrained over this period, this

divergence is surprising and may be related to the ability

of individual radiation codes to simulate heating rates at

lower pressures. 

The zonal mean cooling at 100 hPa and 50 hPa (not

shown) indicates near-zero cooling at all latitudes at 100

hPa and about 0.1 K/decade cooling at 50 hPa in most

models. The GISS model gives considerably higher

cooling at higher latitudes than other models; it reaches

0.5 K/decade at 60°N and S at 100 hPa and about 1

K/decade at the same locations at 50 hPa.

4.4.2.3 THE ROLE OF TRENDS IN STRATOSPHERIC

WATER VAPOR

Observations of increases in stratospheric water

vapor (Section 4.3.2) have led to renewed interest in its

climatic implications (Forster and Shine, 1999, 2002;

Dvortsov and Solomon, 2001; Oinas et al., 2001; Shindell,

2001; C.A. Smith et al., 2001). A major difficulty in comparing estimates of the impact on temperature of stratospheric water vapor changes is in estimating their temporal

and geographical variation. Calculations with idealized

water vapor changes (i.e., 700 ppbv from a background of

6000 ppbv) adopted by Forster and Shine (1999) and Oinas

et al. (2001) are not included here. As noted by Forster

and Shine (2002), it is not so much the idealized nature of

the change that biases these estimates, but more the idealized background water vapor that is adopted. Both Oinas

et al. (2001) and Forster and Shine (2002) demonstrate a

strong dependence of the effect of a given perturbation on

the assumed background concentrations.

–90 –60 –30 0 30 60 90

Latitude

–3

–2.5

–2

–1.5

–1

–0.5

0

0.5 Temperature Trend (K/decade) Berlin GCM

Reading IGCM

Imp.Coll. (2-D)

Reading FDH (NBM)

Reading FDH (ECHAM)

Reading FDH (Zhong)

GFDL GCM

SSU15X Observed Trend (79-97)

50 hPa

Figure 4-25. Global and zonal mean temperature trends at about 50 hPa for models using imposed heightresolved ozone trends. The observations from SSU channel 15X are also shown. The 2s error bars in the

observations are included. The peak of the SSU channel 15X weighting function is at about 46 hPa. Acronyms

as defined in text and in Figure 4-24.
GLOBAL OZONE

4.35

Figure 4-26 shows the vertical profile of the modeled global and annual mean temperature trend, derived

for water vapor changes specified ideally (e.g., constant

percentage global increases) or derived from HALOE

observations. Much of the divergence between the models

is likely due to differences in the adopted water vapor

trend; this uncertainty is severe. In the mid-to-upper stratosphere the models all produce a cooling of about 0.2

K/decade, but at lower levels there is significant divergence. The latitudinal structure of zonal mean trends at

100 and 50 hPa (not shown) reflects this divergence; furthermore, at mid-to-high latitudes water vapor increases

could be causing cooling of order 0.5 K/decade.

4.4.2.4 SYNOPSIS OF SIMULATED TEMPERATURE

TRENDS

Figure 4-27 consolidates the above results. An

average profile of cooling has been generated from the

imposed observed ozone change calculations and WMGG

change experiments. For water vapor, the Imperial College

HALOE IGCM trends are used, because these are based on

more global observations, albeit for quite short periods.

Summing these three components generates the “total” trend.

At 1 hPa, ozone change is the dominant contributor to cooling (1.0 K/decade), with WMGGs at 0.7

K/decade and water vapor much smaller at 0.2 K/decade.

Throughout the upper stratosphere, the total model

cooling is within the 2s error bars of the observed cooling.

Ozone depletion and increases in WMGGs contribute

roughly equally to the temperature trends in this region,

although the ozone contribution is more peaked at the

stratopause.

At 10 hPa, WMGGs are the dominant cooling

mechanism (0.4 K/decade), followed by water vapor (0.2

K/decade) and ozone (less than 0.1 K/decade). There is a

significant disagreement with observations. The total

–1.5 –1 –0.5 0 0.5

Temperature Trend K/decade

0.01

0.10

1.00

10.00

100.00 Pressure (hPa)

NOAA Aeronomy 2-D (1%/year, 80-00)

Imp.Coll. 2-D HALOE (92-99)

Imp.Coll IGCM HALOE (92-99)

Reading FDH (NBM) (0.05ppmv/year, 80-00)

SSU/MSU Observed Trends (79-97)

Figure 4-26. Global and annual mean temperature trends for models using imposed water vapor trends. The

observations from MSU and SSU are also shown. The 2s error bars in the observations are included; the vertical bars are intended to give the approximate altitude range sensed by the particular satellite channel.

Acronyms as defined in text and in Figure 4-24.
GLOBAL OZONE

4.36

model cooling at 5 and 6 hPa is outside the 2s errors bars

of the two SSU channels observing at those altitudes, and

indeed, WMGGs alone are toward the edges of these error

bars. A similar disagreement was found in the coupled

model runs (not shown).

The source of this discrepancy is not clear. If the

minimum cooling from any of the contributing models is

used (ozone change (zero cooling); WMGGs (0.45

K/decade); and water vapor (less than 0.1 K/decade)), then

there would at least be agreement within the error bars at

6 hPa. This would require the water vapor trend for the

HALOE period to be unrepresentative of the long-term

trend, but it is also possible that there are inaccuracies in

the observed ozone trend or the observed temperature

trends at these levels.

At 100 hPa, ozone loss is again the dominant

cooling mechanism (about 0.3 K/decade) followed by

water vapor (0.1 K/decade) and WMGGs (near zero

cooling). The net cooling is within the 2s error bars of

the observations. At 50 hPa, the cooling due to ozone

alone is not able to explain the observed trend, and a significant contribution from stratospheric water vapor

appears to be needed to achieve agreement.

Considering next the zonal average trends at 50

hPa, in the tropics at least one of the GCMs with imposed

ozone trends (Figure 4-25) is within the 2s error bars of

the observations; however, in midlatitudes of both hemispheres, there is a substantial difference between the

ozone-induced trend and observations. At 40°N, the gap

between the modeled cooling and the observed error bar

is about 0.5 K/decade. This difference is more than can

be explained by WMGGs alone (most models report a

cooling of about 0.2 K/decade, although one model reports

0.4 K/decade) and would need a cooling of about 0.3

–3 –2.5 –2 –1.5 –1 –0.5 0 0.5

Temperature Trend K/decade

0.01

0.10

1.00

10.00

100.00 Pressure (hPa)

"Well-mixed" greenhouse gas changes

Observed ozone changes

Water vapor changes

TOTAL

SSU/MSU Observed (79-97)

Figure 4-27. Global and annual mean temperature trends from an approximate average of the model results

for the imposed height-resolved ozone trends (Figure 4-24) and greenhouse gases; the water vapor results

are the trends derived from the HALOE results in Figure 4-26. The observations from MSU and SSU are also

shown. The 2s error bars in the observations are included; the vertical bars are intended to give the approximate altitude range sensed by the particular satellite channel.
GLOBAL OZONE

4.37

K/decade from some other mechanism. Water vapor

trends could cause such a cooling, but it requires trends in

excess of those deduced from HALOE and more consistent with the trends used in the Reading FDH model (50

ppbv/year). Such a trend would also improve agreement

at 50 hPa in the global and annual mean plot (Figure 4-

27). In addition, none of the coupled chemistry models

(not shown) lie within the 2s bars of the observations

between 30° and 50°N. Another possibility at these latitudes is that part of the cooling trend during 1979-1998 is

driven by dynamical causes (this is qualitatively consistent with the observed decreases in NH planetary-wave

drag discussed in Section 4.6.3).

For the zonal average trends at 100 hPa, ozone

depletion alone can explain the observed trend at all latitudes within the uncertainty of the observations. The

WMGG trend is small, and a cooling due to water vapor

of several tenths of a Kelvin per decade could be present

and would improve the midlatitude agreement with the

observations, particularly for those models generating a

small ozone-induced cooling.

In summary, there is encouraging agreement

between the observed and model temperature trends, particularly in the upper stratosphere. But the degree of agreement is not always good, and the explanations required to

improve agreement are not always consistent. In the global

annual mean, one way to achieve agreement near 10 hPa

would be for the water vapor contribution to be small, and

yet at 50 hPa at midlatitudes in the annual zonal means, a

significant water vapor contribution is one explanation that

would improve agreement with observations.

4.5 CHEMICAL INFLUENCE ON PAST

CHANGES IN OZONE

This section discusses the role of chemistry in

causing the observed long-term changes in global ozone.

Section 4.5.1 discusses updates to our understanding of

the important gas-phase and heterogeneous chemical

processes. Section 4.5.2 describes the modeling tools that

are used to quantify the role of these chemical processes

on ozone changes and their key uncertainties. Section

4.5.3 then assesses the role of chemistry in driving extrapolar global ozone changes, based on 2-D and 3-D assessment model simulations compared with the observed

ozone changes from Section 4.2.

4.5.1 Chemical Processes that May

Contribute to Ozone Changes

In the photochemically controlled US region above

about 25 km, ozone trends may be driven by trends in the

gases that provide the sources for reactive radicals (such

as N2O, which produces NOx, and H2O, which produces

HOx). Models summarized in WMO (1999) showed that

the dominant driver for ozone trends in this region has

been the observed trends in reactive chlorine (which are

due in turn to the known trends in chlorofluorocarbons).

Observed changes in CH4, H2O, and N2O also have the

potential to contribute to chemical ozone changes, but

their contributions were estimated to have been relatively

small in the past two decades (less than 1% per decade at

40 km). Furthermore, the chemistry of the US is strongly

dependent upon temperature, and therefore it is important

to consider the feedbacks between ozone depletion and

temperature changes in model estimates, and to consider

changes in radiatively important gases (especially CO2

and H2O; see, e.g., Evans et al., 1998).

In the LS and LMS, where the chemical lifetime of

ozone is comparable to or longer than the transport time

scale, both dynamical and chemical processes have the

potential to affect ozone and its trends. In particular,

changes in chemical loss processes may deplete ozone, but

only if they can remove ozone rapidly enough to compete

with the seasonally varying transport that moves ozone

through this region of the atmosphere. Trends in reactive

chlorine may chemically deplete ozone at these altitudes,

and can be particularly effective in combination with reactive bromine. Iodine may influence the background concentrations of ozone near the tropopause, but is unlikely to

be present in large enough abundances to contribute significantly to the observed trends in the midlatitude ozone

column based on the most recent observations.

Ozone in the midlatitude LS and LMS can be

affected both by chemical processes occurring locally and

by chemical changes occurring in other regions, notably

the winter/spring polar vortices, which are then transported to midlatitudes. In this way trends in stratospheric

halogens may drive midlatitude ozone trends either by in

situ midlatitude chemistry, involving heterogeneous

chemistry on liquid sulfate aerosols, or by chlorine activation and/or ozone loss in the polar regions (see Chapter

3) that is then exported to lower latitudes.

4.5.1.1 GAS-PHASE CHEMICAL PROCESSES

Since WMO (1999) there have been many refinements to our understanding of gas-phase reactions

including changes in rate coefficients, identification of

new and potentially important reaction channels, and

reduced uncertainty limits on a host of evaluated rate constants for halogen, HOx, and NOx reactions. The discussion in WMO (1999) was based on data summarized in

the Jet Propulsion Laboratory (JPL) 1997 recommenda-
GLOBAL OZONE

4.38

tion (DeMore et al., 1997). Since then a further recommendation has been published, JPL2000 (Sander et al.,

2000). Here we summarize the main updates (see Table

4-3) and separate the discussion into those updates that

were contained in JPL2000 and those that are more recent.

Whereas the implications for ozone trends of the data in

JPL2000 have been assessed (e.g., Section 4.5.3), the full

implications for the more recent data have not been

determined.

Updates Included in JPL 2000

The discovery of a small hydrogen chloride (HCl)

product channel in the hydroxyl radical + chlorine

monoxide (OH + ClO) reaction reported in WMO (1999)

has been confirmed by further laboratory studies (Lipson

et al., 1999; Wang and Keyser, 2001; Tyndall et al.,

2002). The HCl channel substantially impacts modelobservation comparisons of atmospheric HOCl/HCl,

ClO/HCl, and ClONO2/HCl ratios (Chance et al., 1996;

Dubey et al., 1998). However, because of the relatively

large dispersion in laboratory results (factor of 2 range in

derived rate constant), the magnitude of the HCl channel

remains highly uncertain.

Reactions involving NOx have received substantial

attention, motivated in large part by observations of reactive nitrogen partitioning in the summer Arctic (see discussion in Chapter 3). Recent laboratory measurements

have led to significant revisions in rate coefficients for

O + NO2, OH + NO2 + M (where M is any atmospheric

molecule that collides and dissipates the reaction’s

released energy), and OH + HNO3, especially at the low

end of the temperature range (Sander et al., 2000; Brown

et al., 1999, 2001). The changes in rate coefficients are

found to substantially increase NOx abundance and NOxcatalyzed O3 destruction in the lower stratosphere

(Portmann et al., 1999).

Updates Since JPL 2000

There is continued uncertainty regarding the OH +

NO2 + M reaction, particularly the reconciliation of lowand high-pressure measurements. The reaction has now

been shown to occur via two channels, one producing

HNO3, the other giving rise to pernitrous acid (HOONO).

The laboratory data indicate that HOONO is a relatively

weakly bound molecule (Donahue et al., 2001; Golden

and Smith, 2000) that decomposes back to OH + NO2.

This finding is supported by theoretical quantum chemical calculations (Li and Francisco, 2000). Because of its

weak bond, HOONO will have a much shorter atmospheric lifetime than HNO3, and the fraction of the OH +

NO2 reaction forming HOONO will not effectively

sequester NOx. Inclusion of the HOONO channel in

atmospheric models could potentially affect NOx levels

under high-pressure and/or low-temperature conditions

where the HOONO channel takes on its greatest significance.

Another NOx reservoir, peroxynitric acid

(HO2NO2), has been found to have a photolytic decomposition pathway in the near infrared through excitation to

the first vibrational overtone (Roehl et al., 2002). This

path serves to substantially increase the rate of HOx production under high solar zenith angles and may explain

observed OH behavior at sunrise in the lower stratosphere

(see Chapter 3).

New data on HOx reactions (Herndon et al., 2001;

Nizkorodov et al., 2000) have reduced the uncertainty

associated with the rate coefficients of hydroperoxy radicals plus ozone (HO2 + O3) and OH + O3. Extensive observations of HO2/OH ratios in the lower stratosphere have

provided stringent tests of the measured rate coefficients

and support the results of the newer laboratory studies

(Lanzendorf et al., 2001).

For halogen chemistry, several studies have shed

new light on the kinetics and mechanism of the ClO + HO2

reaction (Nickolaisen et al., 2000; Knight et al., 2000). In

contrast to previous work, the recent studies find a different temperature dependence, which implies a slower

atmospheric cycling through the catalytic mechanism

involving HO2 + ClO and increased importance of longerlived reservoir species ClONO2 and HO2NO2 formed via

reactions of HO2 and ClO with NO2.

Further progress has been made in defining the

chemistry of iodine in the stratosphere. Kinetics data have

been reported for iodine monoxide (IO) reactions with

HO2, bromine monoxide (BrO), and IO (Cronkhite et al.,

1999; Knight and Crowley, 2001; Rowley et al., 2001;

Bloss et al., 2001). The rate coefficient measured for

IO + HO2 is ~30% higher than that used in the calculations of Solomon et al. (1994). This will slightly increase

the (small) calculated impact of iodine species on stratospheric ozone. 

4.5.1.2 MIDLATITUDE LOWER STRATOSPHERIC

HETEROGENEOUS PROCESSES

In the midlatitude LS, heterogeneous reactions

occur mainly on liquid sulfuric acid aerosols. This heterogeneous processing causes the chemical loss of ozone

to be dominated by HOx reactions, with an important contribution from halogen species. Given an accurate knowledge of atmospheric temperature, there are three main

limitations on our ability to quantify the importance of
GLOBAL OZONE

4.39

Table 4-3. Changes in the data on gas-phase reactions relevant to stratospheric ozone photochemistry. Minor changes in some of theother reactions are not listed. Reactions for which the most recent revisions were included in JPL 2000 (Sander et al., 2000), and hence in theAssessment model runs in Section 4.5.2, are indicated by “[JPL 2000]” in the final column.

Reaction Information for 1998 Update Impact Reference

Assessment (WMO, 1999)

OH + NO2 Æ products Few data at lower Refit of p-dependence at Increased NOx/NOy and O3 Sander et al. (2000),

stratospheric T T; decrease k by ~30%; loss rate in 20-30 km region Golden and Smith (2000)

small yield of short-lived region [some updates in JPL 2000]HOONO suggested

O + NO2 Æ NO + O2 Few, and scattered, Measured over larger T range; Decreased NOx catalytic cycle Sander et al. (2000)

data at low T increase k by ~20% at low T between 24 and 36 km [JPL 2000]

OH + HNO3 Æ H O + NO 2 3 Few data at lower T Improved characterization Increased NOx/NOy and O3 Sander et al. (2000),

of p- and T-dependences; loss rate in 20-30 km Brown et al. (1999),

increase k by ~50% in region Brown et al. (2001)lower stratosphere [JPL 2000]

OH + O3 Æ HO2 + O2 ~30% offset between Better definition of T- Improved model fit to Nizkorodov et al. (2000)

various T-dependent dependence; reduced measured HO2/OH

datasets uncertainty

HO2 + O3 Æ OH + 2O2 Low T-dependence Additional low T data; Improved model fit to Herndon et al. (2001)

uncertain, possibly uncertainty reduced measured HO2/OH

nonlinear

Cl + O3 Æ ClO + O2 Scatter in data increases No new data; reanalysis Compensates for change in Sander et al. (2000)

at lower T increases k at low T by ~7% Cl + CH4; no impact on ozone [JPL 2000]

Cl + CH4 Æ HCl + CH3 Data variance due to No new data; reanalysis Compensates for change in Sander et al. (2000)

nonlinear T- increases k at low T by ~7% Cl + O3; no impact on ozone [JPL 2000]

dependence and reduces uncertainty

ClO + HO2 Æ HOCl + O2 Poorly defined T- T-dependence better defined Decreased chlorine Nickolaisen et al. (2000),

dependence, possibly and linear; k ~50% smaller catalytic cycle efficiency Knight et al. (2000)

nonlinear at low T at low T

ClO + ClO + M Æ Large uncertainty at Low T behavior better Increased catalytic Sander et al. (2000),

ClOOCl + M low T defined efficiency at low T Bloss et al. (2001)

HOCl + hn Æ OH + Cl Large uncertainty Third absorption band Increased production rate Sander et al. (2000),

above 380 nm identified; extend database of active chlorine Barnes et al. (1998)

to 420 nm [JPL 2000]

IO + HO2 Æ products No data below 298K Fast reaction; hypoiodous acid Iodine catalytic cycle Knight and Crowley (2001)

(HOI) product observed defined

T, temperature; p, pressure; k, rate constant.
GLOBAL OZONE

4.40

heterogeneous chemistry on midlatitude LS and LMS

ozone: variations of the surface area densities of sulfuric

acid (H2SO4)/H2O aerosols of about a factor of 2 in volcanically unperturbed periods (cf. Figure 4-19); estimated

uncertainties in the reaction probabilities of 15-30% on

cold H2SO4/H2O aerosols and of factors of 2-3 on ice

(Sander et al., 2000); and large uncertainties in how much

air is processed through humid, cold conditions close

to the local tropopause, i.e., how pervasive cold, dilute

H2SO4/H2O aerosols or subvisible cirrus are in the LMS.

Hydrolysis Reactions on Acidic Droplets 

The important hydrolysis reactions of dinitrogen

pentoxide (N2O5) and bromine nitrate (BrONO2) on

H2SO4/H2O aerosols have large reaction probabilities

that are essentially independent of the temperature and

concentration of the droplet solutions. Recently, progress

has been achieved in the quantitative understanding of

BrONO2 hydrolysis. Heterogeneous bromine chemistry

can potentially enhance OH, suppress the NOx:NOy ratio,

and activate HCl (Hanson and Ravishankara, 1995;

Hendricks et al., 1999, 2000). The heterogeneous/homogeneous sequence BrONO2 + H2O HOBr + HNO3,

HOBr + h Br + OH, and OH + HCl H2O + Cl was

quantified in calculations by Lary et al. (1996), who found

that heterogeneous bromine chemistry in the midlatitude

lower stratosphere increased bromine radicals (BrOx),

ClOx, and HOx each by about 20%. Erle et al. (1998) provided observational evidence for the presence of BrONO2

hydrolysis by showing that measurements of BrO and

chlorine dioxide (OClO) during winter outside the polar

vortex (at temperatures above 200 K) agreed better with

photochemical box model calculations when the heterogeneous chemistry of both chlorine and bromine was

taken into account.

Chlorine Activation on Cold Aerosol Droplets

At very low temperatures, H2SO4/H2O droplets in

the stratospheric aerosol layer grow into liquid polar stratospheric clouds (PSCs) caused by uptake of HNO3 (see

Chapter 3). At about 200 K, i.e., several degrees above

the formation temperature of liquid PSCs, heterogeneous

chlorine activation starts (Hanson and Ravishankara,

1995) due to increasing HCl and ClONO2 solubilities

(Carslaw et al., 1997) before the onset of major HNO3

uptake. It remains an open issue as to what extent such

low-temperature processes may also affect LS chemistry

at midlatitudes (Michelsen et al., 1999). The requirement

of low temperatures concomitant with high inorganic

chlorine (Cly) is best fulfilled at the midlatitude tropopause, or at high latitudes in winter. Keim et al. (1996)

found enhanced ClO mixing ratios of up to 80 parts per

trillion by volume (pptv) peaking about 600 m above the

local tropopause at 38°N, and interpreted this as a consequence of the heterogeneous reactions ClONO2 + HCl and

ClONO2 + H2O. These processes would be supported by

(1) locally increased water and low temperatures in the

vicinity of the tropopause (i.e., high HCl solubility), and

(2) enhancements in aerosol surface areas due to high

aerosol loading (resulting from the Mt. Pinatubo eruption). The observations of Keim et al. (1996) were corroborated by Borrmann et al. (1997), who also showed a

correlation between ClO and aerosol surface area.

Because a cooling of 3 K enhances the rate of ClONO2 +

HCl by an order of magnitude, the presence of the volcanic aerosols appears to provide an additional enhancement of the rate but is not a necessary requirement.

Currently it is unclear how often situations suitable

for chlorine activation directly above the midlatitude

tropopause occur. The few available observations seem

to provide contradictory results. Regular humidity measurements onboard a commercial airliner within the

Measurement of Ozone and water vapor by Airbus Inservice Aircraft (MOZAIC) project provide evidence that

about 2% of all data above the tropopause show supersaturation with respect to ice (Gierens et al., 1999, 2000).

This would turn aqueous sulfuric acid aerosols into dilute

solutions (<40 wt% H2SO4) highly suited for chlorine activation and should occasionally trigger ice clouds (see

below). A recent analysis by J.B. Smith et al. (2001) of

ER-2-borne H2O and ClO observations in the midlatitudes

suggests relative humidities decrease strongly within 500

m above the local tropopause (with negligible supersaturation with respect to ice). Therefore, these authors classify cases like the one described by Keim et al. (1996) as

exceedingly rare. However, they found 0.8% of measurements in the first kilometer above the local tropopause

indicated the presence of enhanced ClO (5 out approximately 660 data points). This value is not so different

from the 2% of supersaturation specified by Gierens et al.

(1999), and therefore there need not be a contradiction.

Potential Chlorine Activation on Ice Particles 

In addition to cold liquid aerosols, ice particles are

also a potential host of heterogeneous reactions in the midlatitude LS. Although ice clouds have been observed

above the high-latitude tropopause, the different temperature profile at midlatitudes must make the particles less

frequent. Borrmann et al. (1997) identified a cirrus event

near the local tropopause near Bangor, Maine (45°N),

accompanied by ClO mixing ratios ranging between a few

and several tens of pptv in the vicinity of the clouds. On
the basis of this measurement, Solomon et al. (1997)

included additional very thin (subvisible) ice clouds

derived from satellite observations (P.-H. Wang et al.,

1996) in their model and suggested that chlorine activation might occur through heterogeneous chemical reactions on aerosol and cloud particles. Bregman et al. (2002)

incorporated tropopause ice cloud fields in a 3-D model

and found local perturbations in ClO and NOx to be

comparable to the observations of Keim et al. (1996).

However, quantitative differences between these studies

are substantial, and Bregman et al. (2002) concluded that

it would seem unlikely that heterogeneous chemistry on

subvisible clouds could explain the observed negative

ozone trends in the midlatitude lowermost stratosphere.

A three-year record of continuous lidar observations at the Observatoire de Haute-Provence (OHP, 44°N,

6°E) suggests ice cloud occurrence frequencies in the

lowermost stratosphere of 10-20%, with 5% of the cloud

top heights observed at least 1 km above the tropopause

(Goldfarb et al., 2001). However, this study relies on temperature sonde measurements made 120 km east of OHP,

and possible changes in tropopause height over this distance lead to uncertainty in these numbers. Given the contradictory results from the ER-2 and MOZAIC data with

respect to extent of ice saturation above the tropopause

(see above), the existence and occurrence frequency of

ice particles (subvisible cirrus) above the midlatitude

tropopause remain highly uncertain. 

Other Reactions and Surfaces

In the search for other heterogeneous reactions,

alternative reaction pathways and reaction partners have

been suggested. For example, Drdla et al. (1999) suggested that heterogeneous reactions of formaldehyde

(CH2O) + HNO3 and nitrous acid (HONO) + HNO3 (Iraci

and Tolbert, 1997; Longfellow et al., 1998) could cause

an increase in NOx of up to 25% in the LMS. Another

proposed heterogeneous reaction on sulfuric acid aerosol

is carbon monoxide (CO) + HNO3 (Fairbrother et al.,

1997). This reaction could also enhance the NOx:NOy

ratio, particularly in the LS (Lary and Shallcross, 2000).

An assessment of these reactions in long-term modeling

studies requires a reduction of the remaining uncertainties in the reaction probabilities. 

Soot is another surface that could potentially be

important for ozone chemistry in the midlatitude LS.

Modeling studies have suggested that ozone could be

destroyed on the soot surface, forming O2, and that NO2

and HNO3 could react to form HONO, which photolyzes

to yield HOx (Bekki, 1997; Lary et al., 1999). However,

recent laboratory studies and measurements indicate that

soot is likely to have only a negligible impact on ozone

chemistry in the lower stratosphere (Choi and Leu, 1998;

Gerecke et al., 1998; Kalberer et al., 1999; Strawa et al.,

1999; Disselkamp et al., 2000).

4.5.2 Quantifying the Chemical Influence

on Ozone: Modeling Tools

Studies of global O3 changes have used twodimensional (2-D) latitude-height models or, recently,

three-dimensional (3-D) models. These models contain

representations of atmospheric transport, radiation, and

both gas-phase and heterogeneous chemistry. Some of

these processes may be coupled (e.g., chemically integrated ozone being used in the radiation calculation).

Alternatively, some models may use fixed fields for

certain processes (e.g., transport specified from meteorological analyses). Different models make different

approximations, for example, 2-D models approximate

the atmosphere as a zonal mean, whereas 3-D models may

use a simplified chemistry scheme because of computational costs. Box 4-1 summarizes how well components

of these global models are known to reproduce key

processes in the stratosphere.

Recent studies have provided more information on

the ability of 2-D and 3-D chemical models to quantitatively explain the midlatitude ozone changes. Considine

et al. (1999) performed a Monte Carlo study to assess the

effects of uncertainties in photochemical data on calculated ozone trends, and found a 46% model uncertainty at

the 1s level. Similar conclusions were obtained by Fish

and Burton (1997), who inferred a 50% uncertainty in

midlatitude ozone loss rates based upon a quantitative

analysis of uncertainties in photochemical data. Both

studies argued that, given these large model uncertainties,

the discrepancy between the chemically modeled and

observed trend could not be used as a critical test of theory.

This conclusion remains valid based upon current photochemical uncertainties. Hence, comparisons of observations and numerical models in this Assessment continue

to rely upon a broad range of factors in assessing their

merits, such as seasonal, latitudinal, and vertical variations in calculated ozone depletion, as well as model-data

comparisons for other trace gases such as ClO.

4.5.2.1 2-D MODELS

Two-dimensional radiative-dynamical-chemical

models are still a widely used tool for assessing the effects

of long-term (many decade) changes in chemical species

(e.g., CFCs) on stratospheric ozone. They are computationally inexpensive and can be used for many sensitivity

studies. However, these models cannot account for all of

GLOBAL OZONE

4.41
the factors that influence interannual variability or trends

in atmospheric circulation, though they may capture some

of the feedbacks. Also, they are not well suited to the

treatment of polar processes, because, for example, they

cannot account fully for the dynamics of the polar vortex

and its coupling with midlatitudes. For similar reasons

they are not well suited to the treatment of transport

processes close to the midlatitude tropopause in the LMS.

They are probably most reliable in the US and LS, outside of the polar regions (WMO, 1999). Even though the

models represent atmospheric zonal means, allowance

can be made for departures from the zonal mean in, for

example, the temperature field, which is important for

capturing the in situ Cl activation on cold sulfate aerosols

(e.g., Solomon et al., 1998; Jackman et al., 1996).

4.5.2.2 3-D MODELS

Three-dimensional models aim to represent the full

altitude-latitude-longitude variations in the atmosphere.

These models are computationally more expensive than

2-D models, especially when using “full” chemistry. Fullchemistry 3-D models are now used routinely, but generGLOBAL OZONE

4.42

Gas-Phase Chemistry

Most (if not all) stratospheric models use photochemical modules that are based on standard compilations of laboratory

data (e.g., JPL 2000; see Section 4.5.1.1). This procedure ensures that results from different models and different runs

can be readily compared. All “full chemistry” models contain a complete description of what are believed to be the

main species and reactions that control stratospheric ozone. Model-model intercomparisons have shown that, when

constrained by the same photodissociation rates, the gas-phase chemical packages from the assessment-type models

agree very well (Park et al., 1999; Kawa et al., 1999). 

Heterogeneous Chemistry

Modeling the processes that can contribute to changes in ozone in the LS and LMS depends critically on including a

realistic treatment of heterogeneous chemistry in the model. For nonpolar studies, the key parameter is the sulfate

aerosol surface area density (SAD) and composition. Particularly in the LMS below the maximum in aerosol abundance, the SAD is subject to large uncertainties (of the order of 30%; see Section 4.3.1). For the period perturbed by

volcanic eruptions, the inversion of satellite observations to infer aerosol densities below the peak of the aerosol distribution is a source of uncertainty. For diagnostic studies of observed trends, global 2-D and 3-D models can use observations based on satellite data (e.g., SAGE II back to 1984). Outside of this period, global distributions need to be derived

from other data. 

Temperature

Even with a realistic SAD, model studies have demonstrated that temperatures must be known to better than 2 K to simulate accurately the highly temperature-sensitive Cl-activating heterogeneous chemistry that is critical to ozone loss

(see, e.g., Solomon et al., 1998; Tie et al., 1997). Models that use meteorological analyses will almost certainly have

more realistic temperatures than 2-D models or GCMs in which they are calculated. However, in the past, meteorological analyses have been far less accurate than this stringent 2-K criterion in the polar vortices of both hemispheres (see,

e.g., Manney et al., 1996), although recent analyses show much better agreement (Pommereau et al., 2002). In any case,

uncertainties in polar temperatures add to our uncertainties in the ozone depletion in polar regions (see Chapter 3) and

thus in the transport of ozone-depleted air to lower latitudes (van den Broek et al., 2000).

Transport

Accurate assessment of the cause of global ozone changes requires models with a realistic treatment of the slow, meridional Brewer-Dobson circulation and eddy transport. Since WMO (1999), a great deal of work has been carried out on

comparing models with tracer observations using age-of-air diagnostics (Hall et al., 1999; Park et al., 1999). These

studies showed that many 2-D and 3-D models had vertical transport that was too rapid, and hence they underestimated

the mean age in the stratosphere derived from observations (see Figure 4-28, p. 4.44). This implies that the models

would transport CFCs too rapidly, giving insufficient time for the release of Cly, for example, from halogen-containing

source gases. This may adversely affect trend calculations, not only through the incorrect Cly values in the lower stratosphere, but also because of possible errors in the rate of ozone replenishment through transport. However, the effect of

this uncertainty on trend calculations has not been quantified.

Box 4-1. Components of 2-D and 3-D Chemical Models
ally for short periods of a few months, e.g., to study seasonal polar ozone loss. However, recently full-chemistry

3-D models have become more widely used for multiannual

(or decadal) time scale integrations (e.g., Chipperfield,

1999; Rummukainen et al., 1999; Austin et al., 2000).

Three-dimensional models can be classified into

general circulation models (GCMs) and off-line chemical

transport models (CTMs) that use winds specified from

an external source. When these winds (and temperatures)

are taken from meteorological analyses, the model will

have a reasonably realistic representation of the atmosphere on a given day, which is useful for interpreting specific observations. However, CTMs are still dependent

upon the accuracy of the analyzed temperature fields, and

errors of a few degrees may cause substantial errors in

calculated ozone losses. The use of analyzed winds may

lead to horizontal and vertical winds and temperatures

that are not self-consistent. 3-D models (like 2-D models)

are also subject to errors in the treatment of the slow,

meridional circulation (see Figure 4-28). GCMs calculate their own (self-consistent) winds and temperatures,

but they tend to suffer from large uncertainties in the key

parameter of calculated polar temperatures (see Chapter

3). Although GCMs should be able to more accurately

simulate interactions between the stratosphere and the troposphere than 2-D models, they are still subject to uncertainties, raising questions about the accuracy of some

calculated feedbacks to the dynamical fields (see Section

4.6). Furthermore, because 3-D models are relatively new

and expensive to use, the exact nature of these models’

circulations over long time scales has not been extensively

characterized.

4.5.3 Quantifying the Chemical Influence

on Ozone: Model Results 

Numerical models of the stratosphere have

achieved substantial success in simulating many of the

observed features of the ozone layer and its depletion.

Previous Assessments (e.g., WMO, 1995, 1999) and

model intercomparison reports (e.g., Park et al., 1999)

showed that both two-dimensional and three-dimensional

numerical models could reproduce the general behavior

of a wide variety of stratospheric constituents, including

reactive species such as nitrogen oxides, source gases such

as CFCs, and ozone itself. More than three decades of

such comparisons between models and observations, and

between differing models of varying complexity, have

been a cornerstone of current confidence in the ability of

models to represent many key stratospheric processes,

albeit with remaining quantitative uncertainties. Here we

briefly discuss new information on modeled trends and

distributions of gases other than ozone, to gain insight

into the factors that are likely to have affected ozone trends

over the past decade. We then discuss current model simulations of past ozone changes. 

4.5.3.1 EFFECTS OF NOX- AND ClOX-RELATED

PROCESSES

Koike et al. (1994), Van Roozendael et al. (1997),

and Liley et al. (2000) presented observations from New

Zealand and Europe indicating decreases in NO2 abundances following the Mt. Pinatubo eruption (see Figure

4-21). These were in general accord with model calculations and support the view that this volcano introduced

major perturbations to the photochemistry of the global

stratosphere during the period from about 1992 to 1995

(see Section 4.6.6).

A recent study by Froidevaux et al. (2000) examined trends in satellite observations of ClO during the

1990s. Near the 2-hPa level, they showed a positive trend

in ClO, which is likely to be caused in part by the trend in

Cly over this period. However, they also demonstrated

that variations in methane (which play a controlling role

in the partitioning between ClO and HCl) dominated the

changes in observed ClO during this period, as illustrated

in Figure 4-29. The observed US methane changes in the

mid-1990s (Nedoluha et al., 1998b; Randel et al., 1999)

are poorly understood and not included in standard

models, raising important questions for the modeled ClO

and hence for model-calculated ozone trends in the upper

stratosphere.

In the lower stratosphere, the ClO time series presented by Froidevaux et al. (2000) showed enhancements

at 22 and 46 hPa following eruption of Mt. Pinatubo, and

that perturbation was fairly well reproduced by a 2-D

model to which the data were compared. Observations of

the ratio of HCl/Cly (Webster et al., 2000) provide a separate and important test of the long-term responses of chlorine compounds to variations in sulfate aerosols in the

lower stratosphere, since increased ClO is expected to be

associated with decreased HCl at those levels. Webster et

al. (2000) examined balloonborne and spaceborne data

from 1985 to 1994 and argued for an impact on HCl/Cly

under cold high-aerosol conditions that exceeds current

model calculations. However, Dessler et al. (1997) also

studied in situ observations of HCl/Cly near 50 hPa from

aircraft data obtained from 1992/1993 (immediately after

Pinatubo) to 1995/1996, showing evidence for chlorine

activation close to that predicted by chemical model calculations.

Thus these recent studies have at least qualitatively

supported earlier work showing that the observed changes

GLOBAL OZONE

4.43
in the ClO and NOx abundances of the lower stratosphere

were generally matched by models that included the

effects of the Mt. Pinatubo eruption (e.g., Fahey et al.,

1993; Solomon et al., 1996). Taken together, the measurements and models continue to provide support for the

view that heterogeneous reactions on sulfuric acid

aerosols suppress the abundances of NOx species (mainly

through N2O5 and BrONO2 hydrolysis), while enhancing

GLOBAL OZONE

4.44

Figure 4-28. Mean age of air estimated from a range of 2-D and 3-D models. In all panels the shaded region

indicates the range of mean ages of 17 2-D and 3-D models, not including Harvard-3D (dotted line), MONASH1-

3D (heavy solid line), GSFC-2D (heavy dashed line), GSFC-3D (light dashed line), UCI23 (dot-dashed line),

and UIUC-3D (light solid line), which are shown separately. The symbols represent observations: mean age

from in situ CO2 (triangles) and in situ SF6 (diamonds). (a) Latitude profile versus in situ aircraft data. (b)

Vertical balloon profiles at 7°N. (c) Balloon flight of September 1996, at 35°N (diamonds and triangles) and

SF6 whole-air samples from 44°N (asterisks). (d) Balloon flight of June 1997, 65°N (diamonds and triangles)

and whole-air SF6 samples at 68°N inside (asterisks) and outside (crosses) the winter polar vortex. From Hall

et al. (1999).
concentrations of ClO. The net effect of increased sulfuric acid aerosols is an ozone increase above about 25

km, where NOx dominates ozone losses (Hofmann et al.,

1994; Mickley et al., 1997, Solomon et al., 1996), and

decrease in the lower stratosphere, where ClO abundances

control the impact of halogen chemistry on ozone loss.

The changes in stratospheric aerosol content during the

1990s following the massive perturbation of Mt. Pinatubo

should therefore be expected to exert a substantial influence on ozone changes during the past decade. Standard

models generally include the observed changes in aerosol

based upon SAGE observations, as noted (Box 4-1), and

hence incorporate a representation of these chemical

effects.

4.5.3.2 EFFECT OF POLAR VORTEX ON MIDLATITUDES

Recent model-based studies have attempted to

quantify the effect of Arctic polar processing on midlatitude ozone amounts. Knudsen and Groob (2000) used a

seasonal reverse domain filling (RDF) trajectory calculation to study the dilution of ozone-depleted air into

northern midlatitudes. They estimated that approximately

40% of the 6.8% decline in midlatitude ozone observed

by TOMS between 1979 and 1997 could be accounted for

by transport of polar ozone-depleted air into midlatitudes.

Other studies have used tracers mapped in equivalent

latitude, or discriminated between polar and nonpolar air

in 3-D CTMs, to evaluate the connection between polar

and midlatitude ozone loss. From the analysis of several

Arctic winters in the late 1990s, Millard et al. (2002)

showed that the contribution to the seasonal midlatitude

ozone loss from high latitudes was strongly dependent on

the meteorological conditions and the stability of the polar

vortex. They found the largest contribution to midlatitude ozone loss for winter 1999/2000, with half the midlatitude loss originating north of 60°N. A much smaller

contribution was found in the winter 1996/1997, characterized by a strong and pole-centered vortex prior to the

final breakup. Using a different model and set of tracers,

Marchand et al. (2002) found similar results for the winter

1999/2000. Godin et al. (2002) used a high-resolution

model together with ground-based ozone measurements

to study the influence of vortex excursions and polar air

filaments on midlatitude ozone amounts during several

winters, finding the largest impact during 1999/2000 due

to the large Arctic ozone loss that occurred that winter

coupled with several vortex excursions at the end of the

winter.

The accumulated effect of polar vortex loss on midlatitudes has been studied by Chipperfield (1999) using

multiannual simulations of the 1990s in a 3-D CTM with

full chemistry and horizontal resolution of 7.5° ¥ 7.5°.

That study estimated that overall PSC-related depletion

processes in or at the edge of the vortex results in 2-3%

less ozone at 50°N throughout the year. The effect was

larger in the Southern Hemisphere (5% less O3 at 50°S),

reflecting the large losses occurring in the Antarctic. 

The fate of ozone-depleted air subsequent to NH

polar vortex breakup was studied by Piani et al. (2002).

They showed that by the end of June, much of the ozonedepleted air above 420 K has been transported from the

polar region to the subtropics. In contrast below 420 K,

most of ozone-depleted air remains poleward of 55°N,

due to the influence of the upper extension of the tropospheric subtropical jet that provides a transport barrier at

lower levels.

4.5.3.3 EFFECT OF WATER VAPOR TRENDS

Observations suggest a positive trend in stratospheric water vapor during the past several decades, which

can in turn increase HOx concentrations and feed back

into other chemical cycles involving NOx and ClOx.

Solomon (1999) and Kirk-Davidoff et al. (1999) also

noted the strong dependence of heterogeneous chemistry

on water vapor content; the former work emphasized the

lowermost stratosphere and temperature variability

whereas the latter focused on polar regions. The observed

water vapor increases to date substantially exceed the contribution expected from methane oxidation in the lower

stratosphere, raising important questions about the source

GLOBAL OZONE

4.45

0.20

0.30

0.40

0.50

0.60

0.70

0.80

0.90 ClO (ppbv) and CH4 (ppmv)

ClO from MLS

91 92 93 94 95 96 97

30°S-30°N ClO from MLS and CH4 from HALOE at 2 hPa (42 km)

CH4 from HALOE

OJAJOJAJOJAJOJAJOJAJOJAJO

Figure 4-29. Time series of zonal mean abundances

of ClO from Microwave Limb Sounder (MLS) data

(filled circles) and CH4 from Halogen Occultation

Experiment (HALOE) observations (open circles) for

2 hPa (42 km) and averaged over 30°S-30°N. From

Froidevaux et al. (2000).
of the increases (see Section 4.3.2). While standard photochemical models do not include this additional water

vapor, several modeling groups have carried out specific

tests of its likely effects, guided by observations of the

magnitude of the water vapor change. Evans et al. (1998)

showed that the water vapor trend was unlikely to make a

significant contribution to ozone trends in the upper stratosphere (see also WMO, 1999). Dvortsov and Solomon

(2001) showed that the observed trend in water vapor

could have contributed about 1% to the total column and

peak ozone depletions at northern midlatitudes. They

argued that this depletion was due to enhancements in

HOx chemistry and to the coupling between HOx and ClOx

chemistry in the lower stratosphere (which contributed

about 0.2%/decade).

Aircraft emissions can potentially affect the composition of the atmosphere. The Intergovernmental Panel

on Climate Change (IPCC) Special Report on Aviation

and the Global Atmosphere (IPCC, 1999) summarized the

present understanding and concluded that the 1992 subsonic fleet was responsible for total ozone column changes

(at 45°N latitude) of approximately +0.4%, with approximately a quarter of that change occurring in the stratosphere. This impact is likely to increase with future

growth in fleet emissions (see Section 4.8.3).

4.5.3.4 ASSESSMENT MODEL SIMULATIONS OF PAST

OZONE CHANGES

Within the framework of this Assessment, nine

models have performed simulations of past ozone trends

in order to test current understanding. The participating

2-D models, which have all been widely used in scientific

studies of stratospheric ozone, are listed in Table 4-4. For

this Assessment the models were run with the observed

surface halocarbon abundances described in Chapter 1

(see Appendix 4B). The surface values of N2O, CH4, and

CO2 were taken from IPCC (2001). (Note that these model

runs ignore other drivers of chemical change, such as additional H2O trend or aircraft emissions.) The models used

photochemical data from JPL 2000 (Sander et al., 2000),

which is the most recent recommendation available. The

aerosol distribution in all models was specified from a

monthly climatology based on satellite data, similar to

that used in Jackman et al. (1996) and updated by D.B.

Considine (NASA Langley Research Center).

The previous Assessment (WMO, 1999) demonstrated that these models simulate many features of the

state of the ozone layer prior to the onset of major depletion reasonably well. They are able to reproduce the

observed ozone column minimum in the tropics, the interhemispheric differences in the ozone maxima in the extratropics, and the general features of the seasonal cycles of

both hemispheres. One important difference in the model

simulations is that two models (NOCAR and GSFC-INT)

incorporate interactive temperature changes due to

changes in greenhouse gases (GHGs) (particularly CO2)

and in ozone itself, which result in US cooling of ~1

K/decade. This feedback has a significant effect on simulations of past and future ozone changes in the US, as seen

below.

Figure 4-30 shows the annually averaged ozone

profile trends for 1980-2000 at northern and southern midlatitudes as calculated by eight of the 2-D assessment

models. In the photochemically controlled region above

25 km, there is a large variation in the (maximum) modelcalculated trends, which range from around –4 to

–5%/decade (NOCAR, GSFC-INT and OSLO) to –8 to

–10%/decade for the other models, with observations in

the middle (–6 to –7%/decade). The most important difference between the models is the temperature feedbacks

included in the NOCAR and GSFC-INT simulations.

While the non-interactive models overestimate the

observed ozone trends for 40-50 km (probably due in part

to the neglect of stratospheric cooling), the interactive

model trends are significantly smaller than observations

over this region. However, the observed shape of the profile is well reproduced by all the models, and this is controlled by ClO/Cly partitioning. Although there is confidence that halogen chemistry is the main driver for the

observed trend in the upper stratosphere (WMO, 1999),

the model-model differences indicate the important feedback that other changes (especially temperature) have on

modulating the ozone depletion.

At lower altitudes the models underestimate the

observed NH ozone trends over this period. The model

values cluster in the range of –1 to –3%/decade, while the

SAGE and ozonesonde trends near 20 km are approximately –3 and –5%/decade, respectively. One caveat is

that the ozonesonde trends (derived from an eight-station

average) are larger than can be reconciled with the zonal

mean column ozone trends, and are likely an overestimate

due to limited sampling (see Section 4.2.4). However, the

model trends are probably still smaller than observed

changes in this region, and because 70-85% of the column

trends originate below 25 km (Section 4.2.4), these profile underestimates are the main reason why modeled

column ozone changes in NH midlatitude are less than

observations (shown below). There may be several reasons for these differences. Part of the model trend underestimate (of order 1%/decade) is likely to be due to the

neglect of the observed water vapor trend, and uncertainties in kinetics may also contribute. Further, some recent

observations suggest that the bromine content of the lowGLOBAL OZONE

4.46
ermost stratosphere may be larger than estimated in these

models that are based on long-lived bromine sources only

(Harder et al., 2000). The bromine content of the LMS

may be affected by transport of short-lived substances

such as bromoform and BrO to the upper troposphere and

stratosphere (Ko et al., 1997; Dvortsov et al., 1999;

Schauffler et al., 1999; Sturges et al., 2000; Fitzenberger

et al., 2000; Pfeilsticker et al., 2000; and Chapter 2). Even

if there are no trends in these largely natural sources of

Bry, they could enhance anthropogenic ozone depletion

through chemical coupling with Cly trends, especially via

the ozone-destroying reaction of ClO + BrO. A further

reason for the discrepancy in northern midlatitudes may

be related to the formulation of the 2-D models, in that

they do not accurately simulate polar ozone loss and its

transport to midlatitudes. Finally, dynamical contributions could also contribute to past ozone changes over NH

midlatitudes, as discussed in Section 4.6. 

Figure 4-31 shows the observed and calculated

annual mean column ozone trends as a function of latitude for the period 1980-2000. The largest trends occur

at high southern latitudes, associated with the Antarctic

ozone hole. The models show a wide variation in their

ability to reproduce this feature, with most underestimating the observed trend, though the GSFC and GSFCINT models produce very large depletions (which extend

into SH midlatitudes). In the Arctic, the inter-model differences are smaller and match the observations fairly

well near 60°N. The models tend to underestimate the

northern midlatitude column ozone loss, consistent with

the profile comparisons (Figure 4-30) discussed above.

In the tropics, all of the models overestimate the observed

GLOBAL OZONE

4.47

Table 4-4. 2-D models participating in simulations of past ozone trends.

Model Name Institution(s) Investigators Temperatures PSC Scheme Reference

AER AER, Inc., U.S. D.K. Weisenstein, Specified from Based on model T Weisenstein et al. 

C.J. Scott, NCEP analyses (1998)

M.K.W. Ko

MPIC-2D MPIC, Germany C. Brühl, Specified from Based on prescribed Groo et al.

J.U. Groob, CIRA/MAP data T and probability (1998)

P.J. Crutzen function

GSFC NASA Goddard, C.H. Jackman, Specified from Distribution based Fleming et al. 

U.S. E.L. Fleming, UKMO analyses on NCEP (1999)

D.B. Considine temperatures

GSFC-INT NASA Goddard, J.E. Rosenfield, Calculated by PSC from [ob- Rosenfield et al.

U.S. D.B. Considine model; interactive served T minus (1997)

with CO2, O3, and 2 K]

H2O

NOCAR NOAA/NCAR, R. Portmann, Calculated above 10 NCEP T for PSC Portmann et al.

U.S. R.R. Garcia, hPa (interactive (1999)

S. Solomon with CO2, O3,

H2O, and CH4);

NCEP analyses

below 10 hPa

OSLO University of B. Rognerud Specified Based on Stordal et al.

Oslo, Norway climatology (1985)

RIVMRIVM, The G. Velders Climatology No solid PSCs Velders (1995)

Netherlands (noninteractive)

SUNY-SPB SUNY, U.S.; S. Smyshlyaev, Specified from Based on NCEP Smyshlyaev et al. 

St. Petersburg, M. Geller NCEP temperatures (1998)

Russia

ULAQ University of G. Pitari, Specified from Based on model T Pitari and Rizi 

L’Aquila, Italy E. Mancini NCAR dataset (1993)

AER, Atmospheric and Environmental Research, Inc.; MPIC, Max-Planck-Institut für Chemie; RIVM, Rijkinstituut voor Volksgezandheiden en

Milieu; SUNY, State University of New York; CIRA, Committee on Space Research (COSPAR) International Reference Atmosphere; MAP,

Middle Atmosphere Project.
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Figure 4-30. Vertical profiles of ozone trends over the period 1980-2000 from observations compared with

eight 2-D model calculations for statistics for 35°N-60°N (top) and 35°S-60°S (bottom). The observations are

from SAGE I+II (solid red line with 2s error bars) and averaged ozonesondes (dashed red line; data only

available in the NH). The trend units are %/decade calculated with respect to the climatology of each model.
(near-zero) trend, which may be associated with model

overestimates of the mixing between the tropical and midlatitude LS. 

Figure 4-32 shows the seasonal variation of the

modeled column ozone trends compared with the observations, with results separated for NH and SH midlatitude

(35°-60°) and polar (60°-90°) regions. In the Antarctic

region, most models underestimate the amplitude of the

seasonality, because of an underestimation of the springtime depletion. The two models that produce the largest

ozone hole depletions (GSFC and GSFC-INT) substantially overestimate the trend in southern midlatitudes at

that time of year. The other models generally reproduce

the observed SH midlatitude trends (within the error bars)

and its weak seasonality. In the Arctic region, although

the models do not capture the same amplitude of seasonal

variation as the observations, they do fall within the 2s

uncertainty. At northern midlatitudes the models do a reasonable job of reproducing the magnitude and seasonality

of the trend, although the winter-spring maximum is

underestimated in most models.

Figure 4-33 compares the smoothed time series of

the observed ozone depletion with the 2-D model results,

for the near-global average (60°N-60°S) and for NH and

SH midlatitudes. Note that for the 60°N-60°S average

(Figure 4-33, top panel) the observations use the data with

the solar cycle statistically removed (Figure 4-4), because

these 2-D models do not include a parameterization of the

solar cycle. The models all overestimate the decadal

ozone changes over 60°N-60°S (Figure 4-33), because of

the overestimate of tropical trends (Figure 4-31). For the

latitude band of 35°-60°N (Figure 4-33, middle panel),

the 2-D models are in broad agreement with many aspects

of the observations, but slightly underestimate the

decreases. Both the observations and the models display

decreases in northern midlatitude ozone following the El

Chichón eruption in the early 1980s and the Mt. Pinatubo

eruption in the 1990s. However, the magnitudes of the

observed changes are larger than most calculations, and

there is also an observed dip in the late 1980s that is not

simulated by the 2-D models. Chipperfield (1999) noted

the important role of polar processing in driving springtime midlatitude ozone decreases following eruption of

Mt. Pinatubo, which may be linked to the underestimate

seen in some of the 2-D models in that time period. Also,

some of this post-Pinatubo dip may be dynamically

induced (see Section 4.6). Current model results are also

sensitive to changes in photochemical rates when compared with earlier studies (such as Solomon et al., 1996;

Jackman et al., 1996; WMO, 1999). In particular, new

data for NOx reactions (used in the runs presented here)

caused reductions in both the calculated trends and in the

amplitude of the estimated Mt. Pinatubo perturbation,

compared with studies carried out for the previous

Assessment (as shown by Portmann et al., 1999). One

preliminary run of the NOCAR model using the updated

kinetics discussed in Section 4.5.1, but not included in

JPL 2000, yielded extremely small changes in calculated

midlatitude depletion.

Figure 4-33 (bottom panel) shows the observed and

calculated trends for the region of 35°-60°S. The observations suggest an overall change for the 20 years of about

GLOBAL OZONE

4.49

Figure 4-31. Latitudinal profile of annual

mean column ozone trends from the

merged TOMS+SBUV2 satellite dataset

(solid red line) (see Section 4.2) compared

with results from eight 2-D models, for

1980-2000. The trend calculations use a

standard regression analysis, including a

QBO term for the observations but not for

the models. The error bars for the observations indicate 2s uncertainties.
6%, while the model values range from 4% to 10%. Each

of the models displays an ozone decrease in the SH following the eruption of Mt. Pinatubo, approximately similar to the model calculations in the NH, but observations

show a much smaller signal. There are relatively large

differences of about 2% between the various satellite

datasets during that time, implying caution regarding

exact details. The largest short-term ozone perturbations

for SH midlatitudes are the low values in the mid-1980s,

which are not simulated in any model and have not been

explained. Closer inspection of the data reveals that the

unusual values occurred during a short period (1985-

1986), when ozone was low over midlatitudes throughout

the SH (see Figure 4-6). The short duration of the perturbation for 1-2 years argues against a solar irradiancedriven mechanism, or attribution to a recurring process

such as the El Niño-Southern Oscillation (ENSO) or the

QBO. Thus these comparisons show that although SH

decadal-scale trends are reasonably well reproduced by

some present models (with a wide spread related to different treatments of the Antarctic ozone hole), there are

mechanisms for shorter-term interannual variability that

are not well simulated by any 2-D model.

As discussed above, it is only very recently that

three-dimensional models have been used to study multidecadal past changes in ozone. Figure 4-34 shows results

from a 20-year run of the SLIMCAT 3-D chemical transport model (CTM) (Chipperfield, 1999) over the period

GLOBAL OZONE

4.50

Figure 4-32. Seasonal variation of observed and modeled column ozone trends at 35°N-60°N (top left), 35°S60°S (top right), 60°N-90°N (bottom left), and 60°S-90°S (bottom right). Observed trends (red lines) were

derived from the merged satellite data using a seasonally varying regression model, with error bars denoting

2s uncertainty estimates.
1979-1998. The model was forced by European Centre

for Medium-Range Weather Forecasts (ECMWF) meteorological analyses in two separate simulations, one with

time-dependent halogen loading and one with constant

(1979) halogen loading. Although the atmospheric

halogen loading in 1979 itself was well above natural

levels, the difference between the two model experiments

shows the modeled chemical loss relative to 1980. The

3-D CTM results produce more interannual variability (on

the time scale of a few years) than the 2-D models, because

of the use of analyzed winds. In the early 1980s and the

early 1990s, and again in the late 1990s, the 3-D model

with the imposed halogen trend matches much of the

observed decreases in column ozone. However, there are

periods, during the mid-1980s and mid-1990s, when the

model variations differ significantly from the observations. There are still large interannual changes, e.g.,

related to the low SH ozone in the mid-1980s, that have

not been explained.

In summary, the 2-D assessment models generally

perform better in simulating the NH changes than the SH

changes, both in the polar regions and in midlatitudes. It

is likely that the models’ ability to simulate the middle

latitudes realistically is related to correctly reproducing

the polar region. The larger losses in the SH polar region

(where the models show large differences) compared with

the NH make this connection more critical. In spite of

these detailed differences, the overall agreement between

the modeled and observed trends in their vertical and latitudinal profiles and seasonal variation support the view

of WMO (1999) that increases in halogens are a main

driver of the observed changes.

GLOBAL OZONE

4.51

Figure 4-33. Time series of column ozone variations from observations (red lines) and from eight

2-D models for latitude bands of 60°S-60°N (top),

35°N-60°N (middle), and 35°S-60°S (bottom).

Changes are calculated in percent with respect to

1980 values for each time series. The observations

for 60°S-60°N use results from Figure 4-4, with the

solar cycle component removed by statistical regression, whereas the midlatitude observations are from

Figure 4-7, with no explicit solar cycle adjustments.

The error bars on the observations represent the

spread of the different datasets, which provides a

minimum estimate of their uncertainty.
4.6 DYNAMICAL INFLUENCE ON PAST

CHANGES IN OZONE

4.6.1 Introduction

Transport of stratospheric ozone is a key factor

influencing its seasonal and interannual variability. The

large seasonal cycle in column ozone over the extratropics (in both hemispheres), as seen in Figure 4-35, is due

primarily to enhanced transport from the tropical source

region during the winter-spring seasons, and the hemispheric differences in ozone amount are mainly a result of

differential transport (larger in the NH). Figure 4-35 also

shows that the interannual variability in the winter-spring

ozone buildup is greater in the NH than the SH, reflecting

the greater dynamical variability of the NH stratosphere.

Given the observed large interannual and decadal-scale

variability in stratospheric dynamical quantities, it is reasonable to consider a dynamical influence on decadal

ozone trends on both regional and hemispheric scales

(e.g., Hood and Zaff, 1995).

In WMO (1999), as well as in previous Assessments, the possible influence of long-term changes in

GLOBAL OZONE

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trend)

Figure 4-34. Time-series of column

ozone change (% change from 1980)

from the SLIMCAT 3-D CTM compared with observations (as in Figure

4-33) for 35°N-60°N (top) and 35°S60°S (bottom). Results from two

model simulations are shown: with a

halogen trend and with constant

(1979) halogen loading. 
dynamical processes on ozone changes was discussed.

However, the statistical trend analyses presented there did

not include explicit dynamical variables (except for the

QBO and solar cycle) because of difficulties of interpretation. One particular concern is that establishing the

existence of correlations using observations is far from

proving cause and effect; sound mechanistic information

is also needed. In particular, it is necessary to assess the

extent to which the dynamical changes might be a consequence of the ozone changes, rather than a cause. A further problem is that using correlations to attribute

causality assumes that linear relationships can represent

what are, in reality, complex nonlinear interactions.

Although this is a relatively minor issue for periodic signals such as the QBO and solar cycle, it becomes problematic when trying to apply daily or monthly mean

correlations to longer time scales.

Recently there has been interest in trying to understand and diagnose the effects of dynamical changes on

stratospheric ozone over decadal time scales, and there is

now substantially more information available than at the

time of the previous Assessment. The relevance of dynamical changes is underlined by related issues such as

the unexplained long-term trend in stratospheric water

vapor (see Section 4.3.2), and the apparent trend in the

North Atlantic Oscillation (NAO) and other hemisphericscale meteorological phenomena (Graf et al., 1995;

Thompson et al., 2000). Furthermore, the opportunity for

long-term dynamical studies on a global scale has been

facilitated by the recent availability of meteorological

reanalysis products.

The relevant studies may be largely grouped into

two categories; neither addresses the cause of the observed

dynamical changes. The first consists of modeling studies

using observed circulation statistics derived from meteorological analyses to drive the transport in the model.

These have been performed in the presence of constant

chemical forcing to determine the ozone response to the

dynamical forcings alone, and in the presence of a

changing chemical forcing so that the two effects can be

assessed together. The second category consists of statistical studies that combine mechanistic understanding of

the atmosphere with the statistical demonstration of

correlations in time series. For such studies the previous

caveat about interpreting correlations still holds, both in

assessing the results of individual studies and especially

in comparing the results from different studies. A simple

example makes this point clearly. Increases in both

tropopause height and mini-hole frequency have been suggested as being responsible for decreases in ozone over

GLOBAL OZONE

4.53

JFMAMJJASONDJ

Figure 4-35. Seasonal

cycle of total ozone (from the

merged satellite dataset)

averaged over 35°N-60°N

(upper set of curves) and

over 35°S-60°S (lower set of

curves), with each curve

corresponding to a different

year in the period 1979-

2000.
northern midlatitudes. However, since they are intimately

linked (mini-holes occur when the tropopause is exceptionally high), the two influences on ozone cannot be

simply added.

This section on dynamical influences is organized

as follows. An overview of key dynamical processes associated with midlatitude ozone trends is provided in Section

4.6.2. The statistical analysis of observations is discussed

in Section 4.6.3, and the impact of dynamical changes

estimated using 3-D chemical transport modeling is

detailed in Section 4.6.4. The possibility that the dynamical changes might be the result of the ozone changes is

assessed, and the dynamical impact of GHG changes considered, in Section 4.6.5. We note that most of the discussion focuses on the NH because uncertainties in the SH

circulation statistics do not permit reliable calculation of

trends in dynamical forcing (e.g., Randel et al., 2002).

4.6.2 Dynamical Processes that May

Contribute to Ozone Changes

Recent studies of dynamical influence on midlatitude ozone trends have discussed variations and trends in

dynamical quantities related to ozone transport. These

studies have broadly fallen into topics focused on two

processes that although ultimately coupled, are in principle independent:

• Stratospheric planetary-wave drag (PWD), which

drives the stratospheric Brewer-Dobson circulation

and is a principal dynamical influence on ozone above

20 km although also affecting ozone below 20 km.

• Tropospheric circulation, which affects the ozone distribution in the lowermost stratosphere (up to about

20 km) through both planetary-scale and synopticscale disturbances and the location of the tropopause.

Decadal ozone changes induced by these processes

can be produced from natural variability, from changes in

climate resulting from greenhouse gases, or from ozone

depletion itself. The importance of knowing the cause

depends on the issue being addressed. To quantify the

ozone depletion that can be attributed to ozone-depleting

substances such as CFCs and halons, the causes of the

dynamical influences on ozone changes are unimportant,

as long as they are not a result of the chemical ozone depletion itself. However, to make valid predictions of future

ozone levels, it is important to understand the fundamental

causes of the dynamical influence on ozone changes. Thus

it is useful to distinguish between the identification of

dynamical contributions to ozone changes, and their

explanation.

The effects discussed above concern the direct

impact of transport on ozone. In addition to these purely

dynamical effects, there are chemical and radiative feedbacks between temperature and ozone that are both positive in the lower stratosphere. These chemical and radiative feedbacks may amplify or skew the effect of natural

dynamical variability on ozone. It is thus not possible to

cleanly separate “dynamical” and “chemical” contributions to ozone changes, attributing a certain fraction to

each, because they are nonlinearly coupled. Whether the

coupling is a first-order or a second-order effect, and what

the causality of the relationship is between dynamics and

ozone, has to be assessed on a case-by-case basis. 

4.6.2.1 CHANGES IN PLANETARY-WAVE DRAG

The global transport of constituents in the stratosphere can be described in terms of advection by the mean

meridional circulation together with eddy transport

effects; both of these quantities are first-order terms in the

zonal mean continuity equation (Andrews et al., 1987;

Mahlman et al., 1980). The stratospheric meridional circulation, known as the Brewer-Dobson circulation (BDC),

is primarily driven by mechanical forcing arising from

stratospheric (and mesospheric) wave drag, plus seasonal

variations in radiative heating. Likewise, large-scale eddy

transports are linked to dissipating waves, and hence to

stratospheric wave drag, although this linkage is difficult

to quantify. This forcing originates from upwardpropagating, tropospherically generated waves. Unforced

natural variability (i.e., excluding solar variations or volcanic eruptions) arises primarily from dynamics through

variability in wave drag, although model results suggest it

can be affected by chemical-radiative feedbacks that could

change in the presence of anthropogenic forcing (e.g.,

Shindell et al., 1999; Butchart et al., 2000). Wave-drag

variability can be internal to the stratosphere, or can arise

from variability in the tropospheric wave sources. The

amount of wave drag within the stratosphere is related to

the vertical component of the Eliassen-Palm (EP) flux in

the lower stratosphere, and this quantity (also referred to

as the planetary wave driving) is a convenient proxy used

to quantify PWD.

Planetary-wave drag (PWD) is largely restricted to

winter and spring for well-understood reasons (Andrews

et al., 1987), and drives the spring buildup of extratropical column ozone seen in Figure 4-35. The relationship

between PWD and ozone buildup is not simple, but a clear

positive correlation has been found in the interannual variability of NH extratropical PWD and total ozone buildup

in January (Fusco and Salby, 1999; Randel et al., 2002),

as illustrated by Figure 4-36. Such dynamically induced

ozone changes would be expected to be seen mainly

between midwinter and late summer; each year ozone

GLOBAL OZONE

4.54
essentially returns to photochemical control by the fall

(e.g., Chipperfield and Jones (1999) for the polar region),

and there is a correlation between winter buildup and

summer loss (Fusco and Salby, 1999; Randel et al., 2002).

From Figure 4-35 it is evident that this statement applies

more to the NH than to the SH, where the late breakup of

the vortex does not allow sufficient time for ozone to reach

photochemical control before the next year’s buildup

begins. In contrast, NH ozone exhibits less interannual

variability in early fall; indeed the spread of values seen

in Figure 4-35 in early fall largely represents the trend

shown in Figure 4-8. The fact that the effect of PWD on

ozone is seasonal and has essentially no interannual

memory suggests (at least for the NH) that decadal

changes in PWD can be expected to lead to decadal

changes in ozone, all else being equal. For such wavedrag-induced variability, changes in the extratropics

should be mirrored (with a negative sign) in the tropics,

and this compensation is indeed seen in interannual

variability of total ozone (Fusco and Salby, 1999; Randel

et al., 2002). Although the basic physics of the connection between PWD and total ozone is well understood, its

quantification via correlations is at best a crude measure

of this connection, and this limits our ability to attribute

observed changes in total ozone to observed changes in

PWD.

4.6.2.2 EFFECTS OF TROPOSPHERIC CIRCULATION AND

TROPOPAUSE CHANGES

The tropospheric circulation also affects the ozone

distribution, with the direct effects confined to the lowermost stratosphere. The relationship between midlatitude

tropopause height and total ozone is well documented

(e.g., Bojkov et al., 1993), with roughly a 20-DU decrease

for every 1-km increase in tropopause height (Hoinka et

al., 1996), independent of season. Day-to-day changes in

tropopause height at a given location are associated with

the passage of synoptic-scale baroclinic disturbances in

the upper troposphere and lowermost stratosphere, which

affect ozone through the depth of the lowermost stratosphere (e.g., Salby and Callaghan, 1993; Vigliarolo et al.,

2001). Longer-term changes are associated with

planetary-scale circulation anomalies, such as the NH

annular mode (see next section), which elevate the

tropopause at certain longitudes and lower it at others.

However, the effect of tropospheric circulation on total

GLOBAL OZONE

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79 80 81 82 83 84 85 86 87 88 89 90 91 92 93

Year

0

10

20

30

40 JAN OZONE INCREASE (DU)

0

3

6

9

12 JAN EP FLUX

TOMS O

EP Flux

Figure 4-36. Upward Eliassen-Palm (EP) flux at 100 hPa (a measure of the stratospheric planetary wave

driving) during January (in normalized units) averaged over 20°-90°N (dashed), compared with the January

increase in total ozone over the same latitude range (solid). The regression of TOMS ozone with EP flux has

a correlation coefficient of 0.74. Adapted from Fusco and Salby (1999). Randel et al. (2002) showed that the

strong positive correlation extends over 1979-2000.
ozone, and its relation to tropopause height, is much less

well understood than the effect of PWD on total ozone.

In particular, the relationship between tropopause height

and total ozone mentioned above applies to single stations

and is mainly associated with reversible transport; there

is no reason to expect it to apply in the zonal mean, or on

longer time scales (e.g., seasonal or interannual) over

which ozone transport is irreversible.

4.6.2.3 HEMISPHERIC VARIATIONS AND ANNULAR

MODES

Both hemispheres exhibit natural variability in

PWD, manifested in “strong-vortex” and “weak-vortex”

conditions in the stratosphere. The associated modes of

variability evident in statistical analyses show out-ofphase behavior between the polar and midlatitude regions,

which is primarily zonally symmetric (i.e., annular).

These hemispheric patterns are referred to as the NH and

SH annular modes (NAM and SAM). The stratospheric

NAM and SAM are observed to be coupled to similar tropospheric structures, in both observations (Thompson and

Wallace, 1998, 2000) and GCM simulations (Shindell et

al., 1999), and there is a strong observed correlation

between NAM index and both midlatitude tropopause

height and total ozone during the winter-spring period

(Brönnimann et al., 2000). The stratosphere-troposphere

coupling occurs on sufficiently rapid time scales that the

relevant mechanisms must be dynamical, through wavemean-flow interaction (Baldwin and Dunkerton, 1999),

although the specific mechanisms are far from clear (see

Section 4.6.5). In the troposphere the NAM is closely

associated with the North Atlantic Oscillation (NAO), and

some studies have used the NAO as a dynamical proxy

for ozone variability (e.g., Appenzeller et al., 2000).

4.6.3 Statistical Analysis of Observations

4.6.3.1 CHANGES IN PLANETARY-WAVE DRIVING

In the NH, there have been long-term changes in

various meteorological indicators over the past 30 years

or so, which together paint a fairly consistent, albeit

incomplete, picture. The Arctic wintertime vortex has

been getting colder and stronger (Graf et al., 1995; Pawson

and Naujokat, 1999; Thompson et al., 2000), and more

persistent (Waugh et al., 1999). Any dynamically induced

component of these changes requires a weakened meridional circulation, which at least over the last 10 years (the

length of the relevant data record) is consistent with variations in long-lived trace gases in the middle and upper

stratosphere (Nedoluha et al., 1998b; Randel et al., 1999).

A weakened meridional circulation requires a decrease in

stratospheric wave driving, which has been documented

in several studies (Fusco and Salby, 1999; Newman and

Nash, 2000; Zhou et al., 2001b; Randel et al., 2002).

However, it should be cautioned that these results are quite

sensitive to the months and time period considered. For

example, Newman and Nash (2000) report significant

negative trends in wave driving for January-February

(1979-1999), whereas Hu and Tung (2002) do not find the

decrease (over 1968-1998) to be statistically significant

for November-January. Randel et al. (2002) analyzed data

during 1979-2000 for each month separately, and find the

wave driving to increase during November and December

and to decrease during January and February (with only

January near statistical significance) (Figure 4-37). This

seasonal variation is consistent with the Arctic earlywinter warming and late-winter cooling seen over the

same period at 100 hPa (Langematz et al., 2002). A weakened meridional circulation implies a decrease in the NH

extratropical winter-spring buildup of ozone, in both polar

and midlatitudes, through a decrease in ozone transport.

Because of the seasonality of the winter-spring

buildup of ozone, with any interannual memory (in ozone)

being largely erased by the return to photochemical control by early fall (Figure 4-35), there is some justification

for extending interannual correlations to decadal time

scales. This would only be strictly appropriate if the interGLOBAL OZONE

4.56

Figure 4-37. Linear trends in monthly averaged NH

eddy heat flux (v ¢T¢ ) statistics at 100 hPa (proportional to the upward EP flux), calculated for the period

1979-2000. Results for each month were calculated

based on meteorological analyses from the European

Centre for Medium-Range Weather Forecasts

(ECMWF), the NCEP Climate Prediction Center

(CPC), and the NCEP/NCAR reanalyses. Error bars

denote ±2s uncertainty. From Randel et al. (2002).
annual correlations were established in the presence of

constant chemical forcing (e.g., Cly), which is not the case.

With this caveat, the extent of the resulting wave driving

contribution to midlatitude ozone trends in winter-spring

then depends entirely on the extent of the changes in wave

driving for the period in question. Between 1979 and the

period 1990-1993, the extent and statistical significance

of wave driving changes is particularly sensitive to the

last year chosen, because of the especially large variability

in wave driving during 1990-1993 (see Figure 4-36), and

the estimated impact on ozone is accordingly highly variable (Fusco and Salby, 1999; Randel et al., 2002). Over

the longer period 1979-2000, Randel et al. (2002) find

that decreases in wave driving (Figure 4-37) may account

for ~30% of the observed changes to total ozone in the

January-March period.

4.6.3.2 CHANGES IN TROPOSPHERIC CIRCULATION

AND TROPOPAUSE HEIGHT

Observations have shown that the NH extratropical tropopause has generally risen in altitude over recent

decades. Radiosonde measurements over both Europe

and Canada show an increase in altitude of about 300-

600 m over the past 30 years, the exact amount depending

on location (Forster and Tourpali, 2001; Steinbrecht et al.,

2001). Consistent increases are also seen both in the

ECMWF reanalysis for 1979-1993 (Hoinka, 1999) and in

the NCEP-NCAR reanalysis for 1968-1997 (Thompson

et al., 2000); although not all longitudes show an increase,

it is present in the zonal mean. An increase in zonal-mean

tropopause height is qualitatively consistent with a

decrease in stratospheric planetary-wave drag via a weakened meridional circulation (Thuburn and Craig, 2000),

but Steinbrecht et al. (1998) found little correlation

between changes in monthly mean tropopause height and

in lower stratospheric temperatures, arguing against this

mechanism. Rather, they found a very strong correlation

with mid-tropospheric temperatures, suggesting that the

observed tropopause height changes are mainly caused

by changes in tropospheric circulation. Moreover, an

analysis of ozone profile changes over Payerne,

Switzerland (Weiss et al., 2001), found that the tropopause

height changes are correlated with significant ozone

changes throughout the lowermost stratosphere, but not

above, suggesting that they are primarily associated with

the observed higher frequency of subtropical intrusions

due to synoptic-scale disturbances (Reid et al., 2000;

Bojkov and Balis, 2001), rather than with a change in the

meridional circulation. However, because of our poor

understanding of what controls the zonal-mean midlatitude tropopause height, it is not at all clear that the correlations obtained for month-to-month variability can be

extended to decadal time scales. The seasonality characterizing PWD influences does not apply to tropospheric

dynamical processes, which affect ozone year-round.

Notwithstanding these concerns, several studies

have attempted to estimate the dynamical contribution to

total ozone trends by statistical regression against

tropopause height or tropospheric circulation indices that

are correlated with it. Steinbrecht et al. (1998) analyzed

Hohenpeissenberg, Germany, ozonesonde data for 1967-

1997 and found that about 25% of the trend in total ozone

could be attributed to tropopause height changes. Forster

and Tourpali (2001) used 11 ozonesonde stations in

Europe and Canada from 1970 to 1997 and found that

about one-third of the observed ozone change below 20

km could be associated with the change in tropopause

height, which is consistent with the Steinbrecht et al.

(1998) result. Both studies concerned annual-mean

changes, since the correlation between tropopause height

and total ozone exists at all times of the year. Weiss et al.

(2001) studied changes in the ozone profile at Payerne for

1967-2000 and showed that the tropospheric dynamical

influence on the calculated ozone trend is essentially confined to the lowermost stratosphere (below about 20 km),

and its magnitude depends strongly on the period considered (Figure 4-38).

During the winter-spring period, there is a strong

correlation between the NAM and various tropospheric

circulation indices including tropopause height. Thus,

Thompson et al. (2000), Appenzeller et al. (2000), and

Steinbrecht et al. (2001) used NAM-like dynamical

proxies in statistical trend analyses of total ozone during

the winter-spring period and found consistent results to

those found using tropopause height, namely, ~1/3 of the

observed total ozone trend being statistically associated

with the dynamical changes. The relationship between

wintertime total ozone, tropopause height, and midtropospheric temperature is illustrated by Figure 4-39.

Using global dynamical proxies such as the NAM

removes the concern about the statistical correlations

being dominated by local dynamical effects, and

Thompson et al. (2000) showed using hemispheric data

that virtually all the spatial structure of tropopause height

and total ozone trends during January-March for 1979-

1993 is related to the NAM.

Other tropospheric dynamical indicators have also

been used: Hood et al. (1999) used 330-K potential vorticity and meridional wind shear from NCEP-NCAR

reanalyses for 1979-1998 as dynamical proxies and argued

that 40% of the February total ozone trends and 25% of

the March trends could be attributed to dynamical

changes. So there appears to be a consensus between the

GLOBAL OZONE

4.57
different statistical studies that for decadal time scales, a

significant fraction (20-40%) of the observed NH midlatitude total ozone change is associated with changes in

tropospheric circulation. The fact that these different

dynamical proxies give reasonably consistent results is

not very surprising, however, given the strong correlation

between the various meteorological indices.

4.6.3.3 COUPLING ISSUES

An important question to address is whether the

recent trends in wave driving and in tropopause height are

somehow linked, especially since they lead to similar

estimates of the dynamical contribution to total ozone

changes. Given the expected association between

the NAM and PWD at stratospheric altitudes, the

stratosphere-troposphere coupling apparent in the NAM

would then imply a connection between wave driving and

tropopause height. However, Waugh et al. (1999) found

no particular correlation between the NAM index and

either the total wave driving or the persistence of the

Arctic vortex. Furthermore, the tropopause height trends

exist in all seasons, suggesting that although they might

lead to PWD changes during winter-spring, the converse

cannot be true. So while there is a certain self-consistency

between the various dynamical trends, without a clear

mechanism connecting the NAM, PWD, and tropopause

height, any such connection in the context of long-term

trends remains speculative. This question will probably

not be answerable until the origin of the changes in the

individual dynamical quantities has been identified.

Having said that, it would seem unlikely that the statistical effects of wave driving and tropopause height (or

NAM) on ozone could simply be added.

4.6.4 Modeling Studies

Another way to estimate the contribution of dynamical changes to ozone changes is by using CTMs driven

GLOBAL OZONE

4.58

Ozone trend (% per decade)

Winter-spring trends Payerne 1980-1997

Winter-spring trends Payerne 1970-2000

Figure 4-38. Payerne winter-spring ozone sounding

trends for 1980-1997 (top) and for 1970-2000

(bottom). On the basis of a statistical analysis, the

observed trends can partly be explained by dynamical changes (solid area), while the remaining trend

(shaded area) is attributed to anthropogenic (chemical) ozone depletion. From Weiss et al. (2001).

9

1 0

1 1

1 2 t ropopause [km]

67 70 73 76 79 82 85 88 91 94 97 00 03

February of year

320

360

400

440

O3 [DU] 244

248

252

256 5 km temperature [K]

total ozone: -18 DU/decade

tropopause: +150 m/decade

T(5km): +1 K/decade

Figure 4-39. February monthly means of total ozone

(top), tropopause height (middle), and temperature

at 5-km altitude (bottom), at Hohenpeissenberg

(47.8°N, 11.0°E). Updated from Steinbrecht et al.

(2001) through February 2002.
by the observed wind and temperature fields. This has

recently become possible on a multiannual basis because

of the advance in CTM capabilities (Chipperfield, 1999).

Hadjinicolaou et al. (1997), using parameterized ozone

chemistry, and Chipperfield (1999), using full chemistry,

found that the variations in NH midlatitude total ozone in

the early 1990s can be largely accounted for by circulation changes alone, including the aftermath of the Mt.

Pinatubo volcanic eruption (Figure 4-40). There is, however, a chemical feedback, in that colder winters lead to

more polar chemical ozone loss, which is transported to

midlatitudes. Hadjinicolaou et al. (1997) identified significant PSC-induced ozone losses both in the polar region

and at midlatitudes, especially during the cold winters of

1996 and 1997. Chipperfield (1999) identified a longterm 2-3% reduction at 50°N from polar processing, and

about a 1% reduction from direct midlatitude Cl activation reaching 3% in 1992/1993. Nevertheless a contribution to the very low total O3 anomaly seen in 1993, relative to the years 1992 and 1994, was attributed to circulation changes (possibly the result of the eruption of Mt.

Pinatubo itself), as illustrated in Figure 4-40. From a similar integration covering the 20-year period of 1980-1998,

Hadjinicolaou et al. (2002) further argued that at least half

of the observed NH midlatitude trend from December

through February during this period could be accounted

for by dynamical forcing (Figure 4-41), and that the modeled total ozone was strongly correlated with both the

NAM and stratospheric wave driving. Hadjinicolaou et

al. (2002) furthermore found that the latitudinal and seasonal characteristics of the dynamically induced modeled

ozone decrease were consistent with observations,

GLOBAL OZONE

4.59

Year

280

300

320

340

360 Dobson Units

92-97

92-93

93-94

92 93 94 95 96 97

Year

-30

-25

-20

-15

-10

-5

0 % Difference

77.5° N

47.5° N

92 93 94 95 96 97

Model O Column 47.5oN 1992-1997 run

Model O Column (PSC) – (no PSC)% 1991-1997 run

Figure 4-40. Results from

a 3-D CTM with parameterized O3 chemistry. Top

panel: Zonal mean

column ozone (DU) at

47.5°N for three experiments: annually varying

meteorological forcing,

repeated 1992/1993

forcing, and repeated

1993/1994 forcing. The

interannual variability in

ozone is controlled by the

meteorological variability.

Bottom panel: Difference

in zonal mean column O3

at 47.5°N and 77.5°N

between two runs (with

annually varying meteorology): one run included

the effects of PSCs and

the other did not. The

colder winters show a

larger difference. From

Hadjinicolaou et al.

(1997).
although the vertical structure was not particularly close.

Results from the parameterized ozone runs of

Hadjinicolaou et al. (2002) can be compared with the fullchemistry 3-D simulations shown in Figure 4-34, which

use constant halogen loading. This shows that dynamical

variations do produce much of the short-term variability,

but the largest contribution to the modeled trend with timedependent halogen loading is due to halogen chemistry,

rather than dynamics. It should be noted that the CTM

results are further from observations in the SH than in the

NH, and in particular do not match the decadal-scale variability well. This is perhaps because of the poorer quality

of the analyzed winds and temperatures in the SH.

A complementary approach to that of CTMs is

using stratosphere-only dynamical-chemical models with

explicit representation of planetary waves, driven by

observed forcing at the tropopause. This approach was

used by Kinnersley and Tung (1998) to study Arctic ozone,

and by Fusco and Salby (1999) to study midlatitude ozone.

Both studies confirmed the large effect of interannual variations in wave driving on total ozone. Fusco and Salby

(1999) found results consistent with those of

Hadjinicolaou et al. (1997), including the latitudinal

structure of the ozone trends.

4.6.5 Cause of the Dynamical Changes

The question then arises as to why the dynamical

changes have occurred. This can ultimately only be

answered with GCMs. The first question to resolve is

whether chemical ozone depletion might somehow have

caused the circulation changes. There has been significant attention paid to stratospheric cooling induced by

ozone losses (see Section 4.4), but considerably less to

associated dynamical feedbacks. 

With regard to stratospheric PWD, there is no clear

mechanism that would predict ozone loss leading to a

decrease in PWD. In contrast to the effect of PWD on

ozone, which is clear (if difficult to accurately quantify),

the effect of the zonal-wind and temperature structure on

PWD is not well understood. Chen and Robinson (1992)

and Limpasuvan and Hartmann (2000) have argued that

stronger vertical shear of the zonal wind in high latitudes

reduces the strength of stratospheric PWD (although they

disagree on the details of the mechanism), whereas Hu

and Tung (2002) argue for precisely the opposite effect.

These studies appeal to different physical mechanisms

than does the study of Shindell et al. (1999), which focuses

on the zonal wind (not its vertical shear) in subtropical

(not high) latitudes. In any case, the seasonality of the

relation between NH PWD and ozone means that springtime ozone depletion cannot be used to explain reduced

wintertime PWD. Although Langematz et al. (2002) find

a reduction in stratospheric PWD from their imposed

ozone changes, this reduction is not statistically significant and only amounts to about one-fourth of the observed

PWD changes. Thus although a contribution from the

dynamical feedback from ozone changes cannot be

excluded, it appears that other factors must contribute to

changes in PWD.

With regard to tropospheric circulation changes, it

is more difficult to draw conclusions, partly because of

our more limited understanding of the tropopause region.

GLOBAL OZONE

4.60

80 82 84 86 88 90 92 94 96 98

Year

-6

-4

-2

0

2

4 Ozone Anomaly (%) Model no PSC chemistry

Model with PSC chemistry

SBUV data

TOMS data

Figure 4-41. Anomaly of

column ozone at 30°N60°N for 1980-1998 from

3-D model simulations with

parameterized ozone

chemistry compared with

satellite data. The chemical forcing is held constant

in time, so the modeled

trend in ozone is dynamically induced. From

Hadjinicolaou et al. (2002).
The observed lower stratospheric cooling attributable to

ozone decreases would be expected to raise tropopause

height on radiative grounds alone. Although this direct

effect has not been quantified, Thuburn and Craig (2000)

removed all the stratospheric ozone in an idealized GCM

experiment and found the extratropical tropopause rose

by only 1 km, which suggests that the observed ozone

losses of several percent could not account for a rise of

300-600 m. Furthermore, as noted earlier, the observed

increase in tropopause height appears to be statistically

associated with tropospheric warming, rather than with

stratospheric temperature changes (Steinbrecht et al.,

1998). Recent GCM studies confirm this: Langematz et

al. (2002) found no significant change in mid-tropospheric

circulation (including the NAM) arising from imposed

ozone changes; Shindell et al. (2001) found an increase in

the NAM index, but not of sufficient magnitude to account

for the observed change; Santer et al. (2002) found that

observed ozone changes had a small positive (but statistically insignificant) effect on tropopause height changes

over the past 20 years. Thus, it appears that although the

ozone decreases may have led to changes in tropospheric

circulation (including tropopause height) that are qualitatively consistent with the observed changes, they are of

insufficient magnitude, implying that other factors must

contribute. However, it should be noted that most GCMs

cannot resolve the observed tropopause height changes,

so they must be inferred by interpolation (see Santer et

al., 2002), and furthermore the observed ozone changes

are not well quantified close to the tropopause.

The subsequent question is whether the circulation

changes might be the result of GHG-induced climate

change. On the basis of radiative balance, the expected

tropospheric warming and stratospheric cooling should

lead to an increase in tropopause height at midlatitudes,

and to a strengthened meridional temperature gradient in

the upper troposphere/lower stratosphere. The GCM

study of Santer et al. (2002) showed a mean extratropical

tropopause pressure decrease of 3 hPa/decade resulting

from changes in WMGGs; this corresponds to an extratropical tropopause height increase of about 120

m/decade, which is roughly consistent with observed

changes over recent decades. Perlwitz and Graf (1995)

proposed that these changes would act to weaken NH

stratospheric PWD, and Shindell et al. (1999, 2001)

indeed found such a dynamical feedback in their GCM

study of the response of the stratosphere to GHG forcing.

However, other GCM studies come to the opposite conclusion; for example, Butchart and Scaife (2001), Schnadt

et al. (2002), and Langematz et al. (2002) found that GHGinduced changes lead instead to stronger PWD and a more

intense Brewer-Dobson circulation (BDC). Caution is

also warranted because the trend in the NAM index seems

to have dropped substantially in the last 5 years or so, and

the recent behavior may well be a manifestation of natural variability. As Hartmann et al. (2000) emphasize, the

fact that the NAM is a mode of atmospheric variability

means that it can be affected by even a relatively small

external forcing, and its response may be highly sensitive.

It also makes secular changes difficult to separate from

natural low-frequency climate variability, because they

bear the same fingerprint.

4.6.6 Effects of Mt. Pinatubo Eruption

The volcanic eruption of Mt. Pinatubo in June 1991

produced an increase in stratospheric aerosols that persisted for several years (e.g., Figure 4-18) and that significantly perturbed stratospheric circulation and chemistry.

There was a decrease in tropical ozone for several months

following the eruption, attributed to a transient increase

in tropical upwelling (Schoeberl et al., 1993). Large ozone

decreases were observed in NH extratropics for several

years following the eruption, primarily during the winterspring seasons (largest in 1992/1993), as documented in

ozonesonde and satellite data (Kerr et al., 1993; Hofmann

et al., 1994; Gleason et al., 1993). Smoothed time series

of NH midlatitude ozone anomalies suggest the Mt.

Pinatubo effect persisted during approximately 1992-1996

(Figure 4-7). However, although the Mt. Pinatubo aerosol

was transported into the extratropics of both hemispheres

and covered much of the globe by early 1992, column and

profile ozone measurements in SH midlatitudes (Figures

4-6 and 4-15) do not show a pronounced ozone decrease

comparable to that seen in the NH. This lack of a SH Mt.

Pinatubo signal in ozone is curious, in light of the 2-D

model simulations that suggest a chemical signal of comparable magnitude between hemispheres (Figure 4-33).

The NH extratropical ozone losses following eruption of Mt. Pinatubo have been characterized as primarily a

chemical depletion in several modeling studies (Brasseur

and Granier, 1992; Tie et al., 1994; Kinnison et al., 1994;

Rosenfield et al., 1997) and in WMO (1995, 1999). The

evidence for a chemical mechanism is provided by observations of increases in stratospheric HNO3 and decreases

in NO2 following the eruption (e.g., Koike et al. (1994); see

also Figure 4-21), enhanced ClO in the lower stratosphere

(Fahey et al., 1993; Froidevaux et al., 2000), together with

the observed vertical profile of ozone changes (decreases

in the lower stratosphere over 12-22 km, and increases

above 24 km; Hofmann et al. (1994)). The lack of a corresponding Mt. Pinatubo ozone signal in SH midlatitudes is

even more puzzling in light of the large perturbations in

stratospheric NO2 that were observed at Lauder, New

GLOBAL OZONE

4.61
Zealand (Figure 4-21). These NO2 anomalies are evidence

of anomalous chemistry acting on the Mt. Pinatubo aerosol

in SH midlatitudes, which would be expected to lead to

increased ClO and associated ozone losses.

The model simulations of Hadjinicolaou et al.

(1997) called into question a purely chemical mechanism

for NH midlatitude ozone loss, because they simulated a

post-Mt. Pinatubo decrease in a CTM based solely on analyzed meteorological fields (Figures 4-40 and 4-41).

Results in the 3-D CTM calculations shown in Figure 4-

34 suggest a contribution of roughly 50% from chemistry

and 50% from dynamics for the NH midlatitude ozone

decrease, consistent with the recent coupled chemistry

GCM study of Al-Saadi et al. (2001). The dynamical contribution to midlatitude ozone loss arises from reduced

transport, and the relation to the volcanic aerosols has

been studied in several idealized models. Observations

show that the volcanic aerosols act to warm the tropical

stratosphere (Figure 4-22), persisting for several years.

Robock and Mao (1992), Graf et al. (1994), and Kirchner

et al. (1999) argued that the change in the zonal winds

implied by thermal wind balance would then reduce the

PWD in the NH winter, essentially putting the atmosphere

into a high-NAM-index state. This dynamical feedback

is expected to operate only in the NH during winter-spring,

because of the strength and sensitivity of PWD, and to be

much weaker in the SH. The decrease in PWD leads to a

weaker BDC during several NH winters following the

eruption, causing less midlatitude ozone from transport

effects, and also colder temperatures in the lower stratosphere (which can enhance chemical loss preferentially in

the NH). The overall effect on ozone would be a combination of dynamical and chemical effects, with an interhemispheric asymmetry. Although these model results

are suggestive, the causal relationships are difficult to

quantify in observations because of natural variability of

the winter stratosphere, and quantified understanding of

the global response to the eruption of Mt. Pinatubo is an

ongoing research topic.

4.6.7 Conclusions

There is now evidence based on both empirical/-

observational and modeling studies that a sizable fraction

of the trends in total column ozone reported over northern

midlatitudes during the past 20 or so years has occurred as

a result of decadal changes in atmospheric dynamics. It is

hard to quantify this fraction accurately, partly because of

the large uncertainty (due mainly to interannual climate

noise) in both the observed ozone changes and the estimated dynamical changes, reflected in the sensitivity of

the trends to the length of the record, and partly because

the effects are not really separable because of nonlinear

feedbacks. There is no doubt that this has occurred on a

regional basis, with, for example, the reported ozone trends

over mainland Europe being significantly enhanced by

long-term changes in dynamics. There is now much

stronger evidence than at the time of the previous Assessment (WMO, 1999) that the same is true for the Northern

Hemisphere as a whole. The available evidence suggests

that the dynamical feedbacks from ozone changes (whether

chemical or dynamical) probably act in the same sense as

the observed changes, but are of insufficient magnitude to

account for them. The cause of the dynamical changes is,

at the present time, unknown; neither greenhouse-gas

forcing nor natural variability can be excluded.

4.7 IMPLICATIONS FOR CLIMATE OF

UPDATED STRATOSPHERIC OZONE

CHANGES

This section presents a brief update of the understanding of the impact of changes in stratospheric ozone

on climate. IPCC (2001) presented a detailed review of

recent work on the climate impact of tropospheric ozone

changes; this is not a topic of this section.

IPCC (2001) gave a central estimate of –0.15 ± 0.1

W m-2 for the radiative forcing due to stratospheric ozone

loss over the period 1979-1997 and included an extensive

discussion of the forcing/response relationship for stratospheric ozone changes. Forster et al. (2001) computed

the ozone radiative forcing using the Randel and Wu

(1999) ozone trends and three different radiative transfer

schemes. They estimate a forcing of between –0.08 and

–0.15 W m-2, which is on the low side of the IPCC estimate. Hansen et al. (2000) and Hansen and Sato (2001)

estimate a forcing of –0.1 ± 0.1 W m-2; this more recent

estimate is less negative than their previous estimate of

–0.2 W m-2 because the more recent ozone trend estimates

they use show more middle-stratospheric ozone loss than

they had used in their earlier estimates. Schnadt et al.

(2002), using ozone trends derived from their coupled

chemistry-climate model, estimate a 1980-1990 forcing

of –0.02 W m-2; this smaller value may be related to that

model’s underestimation of the Northern Hemisphere

midlatitude ozone loss. Bengtsson et al. (1999) performed

time-dependent GCM experiments including observed

changes in stratospheric ozone for the period 1979-1997.

They found that the ozone reductions in that period lead

to a clear cooling, not only in the lower stratosphere, but

also in the troposphere. In their model the relative cooling

effect was larger in the upper than in the lower troposphere. They point out that this reduces the upper tropospheric warming by the well-mixed greenhouse gases,

GLOBAL OZONE

4.62
resulting in an upper tropospheric temperature change in

better agreement with observations. 

Over the past 20 years, stratospheric ozone changes

may have offset about 20% of the forcing due to wellmixed greenhouse gases, consistent with the values

reported in WMO (1999); because the ozone change has

been concentrated in this period, the offset since the 19th

century is much smaller, about 5%.

An area of recent interest is the degree to which

radiative forcing is a useful predictor of climate change;

earlier work was discussed in Section 10.3 of WMO

(1999). A key unresolved issue is whether the surface

temperature response to a given ozone radiative forcing is

the same as that due to an identical forcing from, for

example, changes in carbon dioxide. If we take r to be

the ratio of the global-mean surface temperature response

for a global-mean ozone (or any other) radiative forcing

to the global-mean surface temperature response for the

same global-mean radiative forcing but due to carbon

dioxide, then, as discussed by Shine (2000), there is little

agreement among available GCM results. Hansen et al.

(1997) imposed stratospheric and tropospheric ozone

changes on their model and found a value of r of about

0.7 to 0.8 (for simulations including cloud feedbacks).

Christiansen (1999), using idealized ozone changes, found

r to be near unity for lower stratospheric ozone changes

but 1.4 for upper stratospheric ozone changes. Forster

and Shine (1999), using observed stratospheric ozone

changes, found r to be 1.4. More recently, Stuber et al.

(2001) obtained a value of r of 1.8 for idealized stratospheric ozone changes, as a result of a strong stratospheric

water vapor feedback in their model. Joshi et al. (2002)

provide a comparison of three different GCMs (including

that of Stuber et al. (2001)) using the Stuber et al. (2001)

perturbation. They found that r varies from 1.25 to 1.8,

but interestingly for all three models, the lower stratospheric ozone perturbation generated the largest value of

r amongst a range of other climate perturbations.

Hence there remains a lack of consensus as to

whether the climate is more or less sensitive to a given

radiative forcing due to stratospheric ozone changes than

it is to the same radiative forcing due to carbon dioxide

changes, but there is now some indication that it may be

more sensitive. This will act to modestly enhance the relative importance of stratospheric ozone changes on climate, relative to what would be inferred from radiative

forcing alone.

4.8 FUTURE CHANGES IN OZONE

Following the controls on halogen-containing

source gases due to the Montreal Protocol (and its

Amendments), the stratospheric loading of chlorine and

bromine should decrease in the coming decades (see

Chapter 1). Accordingly, over the time scale of the next

50 years, the stratospheric ozone layer is expected to

“recover” from the effects of halogen-induced depletion.

However, because of other atmospheric changes, future

ozone levels may differ from those in the past. These other

changes include the radiative effects (stratospheric

cooling) of increased GHGs, the chemical effects of certain GHGs (e.g., CH4), dynamical changes, and changes

to the aerosol loading. There are many ways of defining

the recovery of stratospheric ozone. These include: a

decrease (or reversal) of the rate of ozone decline, an

increase in column ozone (or ozone at some specific

altitude), a return to some past ozone level, or a restabilization to a new “unperturbed” level. Under certain plausible scenarios, it is possible that stratospheric ozone will

not return to 1980 levels and indeed may not stabilize

within the next century.

The detection of recovery will require long time

series of high-quality observations. Weatherhead et al.

(2000) estimated the length of ozone datasets needed to

detect a statistically significant increasing trend in ozone.

Even assuming no external perturbations (e.g., from volcanoes), Weatherhead et al. (2000) estimate at least 15

years of observations will be required. This estimate also

assumes that no interruptions occur in the time series (e.g.,

due to data gaps and instrument changes). These calculations show that the unequivocal detection of recovery will

not be made in the near future and emphasize that longterm observations should be obtained.

The stratospheric 2-D models used to interpret past

changes in ozone in Section 4.5 have also been used to

predict the future evolution of ozone over the time period

2000-2050. WMO (1999) discussed the issue of future

ozone and presented a number of 2-D model calculations.

These calculations focused on the sensitivity of the rate

of recovery to different halocarbon scenarios. In this

Assessment the impact of different halocarbon scenarios

on the decrease in the stratospheric halogen loading is discussed in Chapter 1. Here we take the baseline halocarbon

scenario (Table 4B-2 in Appendix 4B) and use the chemical models to investigate the expected evolution of stratospheric ozone under expected GHG changes, and the

sensitivity of this to changes in CH4, N2O, and CO2 from

these scenarios. (Note that the models do not simulate

possible future changes to tropospheric ozone; see Section

4.8.4.) Table 4B-1 (Appendix 4B) lists the model experiments performed. The two standard GHG scenarios, runs

MA2 and MB2, are taken from IPCC (2001) scenarios A2

and B2, respectively (Table 4B-3 in Appendix 4B). IPCC

(2001) described four scenario families that made difGLOBAL OZONE

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ferent assumptions about economic and social development. These scenarios are all considered to be “equally

likely,” and here we have chosen two to illustrate the different sensitivity of CH4 and N2O changes. The 2-D

model runs assumed constant aerosol loading based on

the actual 1997 values. Six sensitivity runs were performed to investigate ±20% changes in the rate of change

of CH4, N2O, and CO2 from 2000 onward. Finally, the

impact of volcanic eruptions was considered by assuming

episodic eruptions (i.e., a repeat of the 1979-2000 aerosol

loading) and a constant, enhanced loading corresponding

to 1993.

4.8.1 Effect of Decreasing Halogens

Figure 4-42 shows the predicted evolution of global

and midlatitude ozone from the 2-D models for run MA2.

The models generally predict a local minimum in column

ozone in about 2000, followed by a steady increase. There

is, however, a large spread in the predicted time at which

global ozone returns to 1980 levels, ranging from 2025

until after 2050. The figure can be compared with Figure

12-6 from WMO (1999), where the model runs were based

on a “maximum allowed production” halocarbon scenario,

and none of the models predicted a return to the 1980

global ozone level before the end of the simulations in

2050. Updates to chemical kinetics, especially key NOx

reactions, are probably an important contributor to the

overall more rapid increase in the simulations presented

here. Note that the two interactive models (NOCAR and

GSFC-INT) have the more rapid rates of predicted column

ozone increases in Figure 4-42, due primarily to changes

in the US discussed below. This more rapid ozone

increase due to the effect of increasing CO2 was also

present in the coupled 2-D model studies of Rosenfield et

al. (2002).

Figure 4-43 shows the predicted evolution of ozone

in the upper stratosphere (38-43 km). The two models

that include temperature feedbacks from increasing GHGs

(GSFC-INT and NOCAR) show a significantly faster

increase in ozone. In contrast, the models that use fixed

climatological temperatures show a more gradual

increase, and by 2050 the predicted ozone is still less than

1980 values. An additional run was performed with the

NOCAR model (NOCAR2) that used fixed CO2 after

2000, thereby keeping upper stratospheric temperatures

effectively constant. The difference in temperature

between these two runs in the US was about 6 K by 2050.

Figure 4-43 shows that this NOCAR2 run exhibits a much

slower rate of increase, in line with the other fixedtemperature 2-D models, highlighting the importance of

temperature changes on future US ozone. The cooling

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Figure 4-42. Predicted future evolution of column

ozone averaged over latitude bands 60°S-60°N

(top), 35°N-60°N (middle), and 35°S-60°S (bottom),

from a suite of eight 2-D models. Results are shown

for GHG scenario MA2. The plots also include the

model results and observations of past trends (red

lines) prior to 2000 (see Section 4.5.3).
not only has the direct effect on the rates of chemical reactions, but also can lead to circulation changes that in turn

will affect the US abundance of, for example, CH4 and

NOy. However, the dominant effect is likely to be the

decreased rate of ozone-destroying gas-phase chemical

reactions due to colder temperatures. This is a wellunderstood feedback (e.g., Haigh and Pyle, 1979) that is

captured by the NOCAR and GSFC-INT models.

However, these models underestimate the observed depletion in the late 1990s, and so their predicted increase starts

from too high a value. For the models with fixed temperatures, the predicted increase in CH4 should also tend to

increase O3 near 40 km (by increasing the rate of conversion from ClO to HCl.). Therefore, other factors (e.g., the

increase in NOx, from N2O, or the increase in HOx) are

apparently exerting a dominant influence by 2050.

4.8.2 Influence of Changing Climate Gases

The IPCC scenarios given in Table 4B-3 in

Appendix 4B show a range in the predicted evolution of

CH4, N2O, and CO2 under different assumptions. The

effect of this uncertainty in CH4 and N2O on the future

evolution of ozone has been investigated using one representative model (which uses fixed temperatures). A

decrease (increase) in the rate of CH4 growth increases

(decreases) the time for O3 recovery to 1980 levels.

However, the effect of a 20% change in CH4 is to shift the

recovery by only about 2 years. This uncertainty in the

growth of CH4 essentially covers the range of all IPCC

(2001) scenario families except B1, which predicts a CH4

growth some 85% less than the A2 scenario. In contrast,

a decrease (increase) in the rate of N2O growth decreases

(increases) the time for recovery, but again the effect is

small. For N2O the various IPCC scenarios represent a

variation in N2O growth of +8% to –50% compared with

A2. Therefore, for both gases the modeled direct chemical sensitivity of the rate of recovery to the range in the

IPCC scenarios is smaller than the inter-model differences

(and uncertainties). Uncertainties in CO2 could lead to a

large uncertainty in future O3 through the important feedback discussed in Section 4.8.1.

A recent 2-D model study by Randeniya et al.

(2002) supports and extends the analysis presented here,

as well as that of WMO (1999), on the direct chemical

effects of future N2O and CH4 increases. These authors

based model runs from 2000 to 2100 on two IPCC scenarios with different rates of N2O increase and both an

increase and decrease in CH4. (Note, however, that their

model ignored the effect of temperature feedbacks from

GHG increases.) Figure 4-44 shows the calculated percent annual change in ozone at 45°N for 2000-2100 from

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( ) Figure 4-43. Predicted future evolution of US ozone

(38-43 km), averaged over latitudes 60°S-60°N from

the eight 2-D models shown in Figure 4-42. An additional model run (labeled NOCAR2) used the

NOCAR model but with fixed CO2 after 2000 (with

CO2 set to 1980 values, accounting for the discontinuity near 2000). The models used GHG scenario

MA2. The plots also include the model results and

observations (SAGE I+II; red lines) of past changes

prior to 2000 (see Section 4.5.3).

2000 2020 2040 2060 2080 2100

-9

-8

-7

-6

-5

-4

45 N

A1Fl-750

B1

A1FI Ozone % change from 1979

Year

1980 1985 1990 1995 2000

-6

-4

-2

0

2

TOMS/SBUV MODEL

Figure 4-44. Percent change in annual average

stratospheric ozone at 45°N from 1979 levels for the

period 2000-2100 calculated with a 2-D model and a

range of GHG scenarios. The GHG scenarios are

based on IPCC (2001) scenarios B1 and A1F1, with

a third scenario based on A1F1 but with reduced CH4

emissions (labeled by A1F1-750). The insert shows

modeled and observed ozone changes from 1980 to

2000. From Randeniya et al. (2002, Figure 2).
their study. In agreement with the fixed-temperature

models shown here, that study found that ozone increases

between about 2000 and 2050. However, different

behavior could occur at longer time horizons. Increasing

levels of N2O in the 21st century should be expected to

increase NOx, which will lead to enhanced ozone loss in

the middle stratosphere. Moreover, as chlorine abundances decline, such increasing NOx would no longer form

ClONO2 in the lower stratosphere and could hence lead

to substantial ozone loss. Only when both N2O is assumed

to stop increasing and methane is assumed to decrease in

the latter half of the 21st century does their model predict

a stabilized ozone layer. This study makes clear that the

large perturbations in NOx, ClOx, and HOx that could

occur in conjunction with various possible future scenarios for N2O, CFCs, and CH4 in the latter half of the

21st century may complicate long-term ozone changes,

and could even prevent a return to 1980 levels. However,

as Section 4.8.1 shows, the strong influence of CO2 on

cooling the upper stratosphere means that the coupled

radiative effects of future CO2 increases also need to be

considered, which will tend to increase ozone over this

time period.

Figure 4-45 shows the sensitivity of global ozone

to different aerosol scenarios. Assuming an enhanced,

but constant, aerosol loading causes lower O3 values

throughout the period 2000-2050. The largest difference

occurs early in the 21st century with an additional decrease

of 2%. As chlorine in the atmosphere declines, the additional effect of the enhanced aerosol loading decreases,

and by 2050 the difference compared with the background

aerosol is an additional 1% global loss. The episodic

aerosol scenario also shows this behavior; the transient

ozone depletion associated with a Mt. Pinatubo-size eruption is largest in the early 21st century (additional 2.5%

loss in 2013, compared with additional 1.5% loss in 2034).

4.8.3 Other Stratospheric Influences

4.8.3.1 DYNAMICAL EFFECTS

As discussed in Section 4.6, two dynamical influences appear to have contributed to NH midlatitude ozone

decreases over 1980-2000: a decrease in wave driving and

an increase in tropopause height. These mechanisms likewise have the potential to influence future ozone. If the

past dynamical changes represent natural variability, then

one cannot extrapolate past dynamical trends, and ozone

recovery could be either hastened or delayed by dynamical variability. If, on the other hand, the past dynamical

changes represent the dynamical response to GHGinduced climate change, then one might expect these

changes to increase in magnitude in the future. This would

decrease future ozone levels and delay ozone recovery.

Unfortunately, our ability to predict future dynamical influences is very poor, for two reasons. First, dynamical processes affecting ozone exhibit significant temporal

variability and are highly sensitive to other aspects of the

atmospheric circulation. This means that the GHGinduced signal is inherently difficult to isolate or to represent accurately in GCMs, while the climate noise is relatively high (especially in the NH). Second, coupled chemistry GCMs, which are the tools needed to address this

question, are still in their infancy (see Chapter 3). As discussed in Section 4.6.5, predictions of GHG-induced

dynamical changes by such models do not even agree on

the sign of the changes: some models predict weakened

PWD, which would decrease midlatitude (and polar) ozone

via reduced transport, while others predict strengthened

PWD. Note that these dynamical changes would also

affect the lifetime of stratospheric pollutants, with stronger

PWD leading to a reduced lifetime (Butchart and Scaife,

2001). (This latter effect acts on the several-year time scale

associated with the BDC, in contrast to the effect of PWD

on ozone transport, which is essentially instantaneous.)

Thus, the direct effect of PWD on ozone transport (no

matter what its sign) would be reinforced over decadal time

scales by altered chemical ozone loss arising from PWDinduced change in stratospheric chlorine loading.

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° °

Figure 4-45. Predicted evolution of global column

ozone (60°S-60°N) for one representative model for

the MA2 scenario, which assumes a constant 1997

aerosol loading (solid line) along with results from

sensitivity experiments that assume a constant 1993

aerosol loading (dotted line) and an episodic loading

that repeats the 1979-2000 values (dashed line).
4.8.3.2 FUTURE EFFECTS OF AIRCRAFT

Future subsonic and potential supersonic aviation

have been calculated to impact atmospheric ozone. The

IPCC Special Report on Aviation and the Global

Atmosphere (IPCC, 1999) summarized the present understanding regarding future aviation emission levels and

their effect on ozone. Model calculations were performed

for forecasted (to 2015) and projected (to 2050) increases

in NOx emissions (1.7 Tg in 1992, 4.1Tg in 2015, 7.2 Tg

in 2050) from subsonic aviation. The calculated impacts

increase with fleet emissions, rising from +0.4% to +1.2%

(total column change at 45°N latitude) between 1992 and

2050. The impact is primarily confined to the troposphere,

with approximately 25% of the change occurring in the

stratosphere.

The impact on column ozone of a potential fleet of

supersonic aircraft was estimated by IPCC (1999) and

Kawa et al. (1999). The estimated effect on column ozone

(at 45°N latitude) in 2050 of a combined supersonic and

subsonic fleet is –0.4% relative to no aircraft and –1.5%

relative to a 2050 subsonic-only fleet, with all the

supersonic-induced change occurring in the stratosphere.

For the presumed emission levels of NOx, the predicted

decrease in ozone is dominated by the influence of

aircraft-emitted H2O on odd-hydrogen abundances.

Aircraft impact studies conducted since the IPCC

(1999) report have examined the sensitivity of model predictions to changing model parameterizations and input

parameters. Schoeberl and Morris (2000) employed a trajectory model to evaluate the transport of subsonic and

supersonic aircraft exhaust into the stratosphere. Their

simulations indicate that there is less upward diffusion of

supersonic emissions than found in Kawa et al. (1999)

and that subsonic emissions do not significantly penetrate

the 380-K potential temperature surface (i.e., they are confined to the lowermost stratosphere). The trajectory model

results point to possible excessive diffusion in the global

models. However, the sensitivity of ozone change to the

reduced transport remains to be explored.

The importance of meteorological conditions on

the calculated impacts of supersonic aircraft has been

examined by Rogers et al. (2000). Three-dimensional

model calculations were performed identically to those

used in the IPCC Assessment, except that the model was

forced with an annually varying meteorological analysis.

The varying meteorology results in a calculated O3 perturbation that exhibits interannual variability that is of

similar magnitude (i.e., 50 ppbv) to the dispersion found

among the models used in the IPCC (1999) and Kawa et

al. (1999) assessments. This finding strengthens arguments presented in IPCC and Kawa et al. regarding the

substantial influence of model transport differences on

the ozone impact computations.

The impacts of supersonic aircraft on stratospheric

sulfate and PSC abundances have also received further

scrutiny. Kärcher et al. (2000) have coupled a detailed

analytical formulation for the near-field aerosol plume to

a far-field plume model and a 2-D global aerosol model.

Predicted changes in sulfate aerosol surface area density

due the operation of a supersonic fleet using the detailed

microphysical treatment are similar to the earlier IPCC

estimates. The Kärcher et al. analysis indicates that the

sulfur-to-sulfate conversion efficiency and emission index

of particulate organic matter are the most important

parameters affecting large-scale cloud and chemical

impacts.

The effects of supersonic NOx and H2O emissions

on PSC formation have been examined by Considine et al.

(2000) using a PSC parameterization formulated for inclusion in a global 3-D chemistry and transport model. The

increased levels of polar NOx and H2O arising from the aircraft emissions are found to induce an earlier onset of PSC

formation, denitrification, and dehydration in the Southern

Hemisphere polar vortex. Resulting increases in active

chlorine are calculated to produce approximately 1%

decreases in lower stratospheric vortex O3 concentrations.

It should also be noted that increases in the use of

other types of aircraft or other routes, not considered in

the Assessments discussed above, might lead to significant emissions in the stratosphere. In particular, a large

increase in the use of small “executive” jets or introduction of near-supersonic commercial jets could affect stratospheric ozone. Also, increased use of subsonic polar

routes could have important implications for the stratosphere. The potential effects of these will need to be

assessed in more detail.

4.8.3.3 ROCKET EMISSIONS

Additional modeling and observation results have

been reported on rocket combustion emissions and plume

wake chemistry since the previous Assessment (WMO,

1999), in which it was concluded that stratospheric accumulation of chlorine and alumina exhaust from current

launch activities leads to small (<0.1%) global column

ozone decreases. The new data support this conclusion.

They also reveal new chemical and physical processes

that could influence future estimates of ozone loss associated with rocket launches.

Large (~100%) ozone loss has been observed and

characterized in the local wakes of Delta II (Ross et al.,

2000) and Atlas IIAS (Ross et al., 1999a) rockets powered by a combination of solid propellant (NH4ClO4/Al;

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ammonium perchlorate/aluminum) and liquid propellant

(LOX/kerosene; liquid oxygen/kerosene). The cumulative ozone losses in the Delta II and Atlas IIAS wakes are

comparable with those observed in solid-propellant-only

Titan IV rockets, even though the total solid rocket motor

(SRM) emissions are a factor of 9 and 11 greater for the

Titan IV rocket than for the Delta and Atlas rockets,

respectively.

This apparent nonlinear scaling of ozone loss with

total SRM emissions is not predicted by plume wake

models because the LOX/kerosene emissions are not

expected to destroy ozone locally. To explain the observations, it has been suggested that exhaust products of the

hydrocarbon-fueled engines of the Delta and Atlas rockets,

possibly soot, catalyze ozone depletion. This suggestion

remains to be tested because suitable gas and particle emission data for LOX/hydrocarbon-fueled rocket engines are

not available. In situ particle number density and composition measurements made recently in an enormous aerosol

cloud at 20 km altitude were linked to a Soyuz rocket

launch 12 days earlier (Newman et al., 2001). These

measurements constitute the first stratospheric data on

emissions from a hydrocarbon-fueled rocket.

Danilin et al. (2001a) investigated the global implications of large, local ozone loss due to chlorine and

alumina emissions from the current Space Shuttle

(SRM/LOX-hydrogen) fleet. They modeled the accumulation of ozone losses in two Shuttle wakes (followed by

dilution) and found this immediate wake effect not to be

globally significant relative to the ozone loss on the longer

time scale of plume dispersion and mixing with ambient

background air globally. The relative significance of

ozone loss accumulated in the plume on intermediate time

scales (i.e., greater than 2 days) has not been evaluated.

The impact of alumina particles on ozone depletion is calculated to increase with surface area density

(Danilin et al., 2001b). Particle measurements in SRM

wakes have revealed a trimodal size distribution, in general agreement with rocket plume wake models (Ross et

al., 1999b). However, the fraction of particle mass contained in the smallest (i.e., submicron) size modes differs

by more than a factor of 5 between data obtained from

large Titan IV and Shuttle SRMs (Ross et al., 1999b) and

from a small Athena II SRM (Schmid et al., 2002). The

large variance may reflect real differences between the

emissions of large and small SRMs or differences in measurement techniques. In either case, because particle surface area density is most influenced by the submicron particle modes, the variance between the results of Ross et

al. and Schmid et al. adds significant uncertainty to the

predicted global impact of SRM emissions. Assuming the

Schmid et al. particle distribution with its larger number

of submicron particles for the larger SRMs would result

in the ozone impacts from emitted alumina and chlorine

being roughly comparable. Alternatively, assuming the

Schmid et al. distribution only for small SRMs would

imply that ozone loss relative to total SRM emission varies

inversely with SRM size, a feature not included in models.

The effect of plume-wake processing on the reactivity of rocket-emitted alumina particles and the associated ozone loss has yet to be fully explored. The reactivity depends strongly on the composition of the wake

particles. Some insight into the composition of rocket

particles has been gained from measurements of H2O

vapor and CO2 in the lower stratospheric wake of an

Athena II SRM (Gates et al., 2002). These data show that

a fraction of the H2O emissions condensed onto particles

and remained condensed for more than 40 minutes after

launch, long beyond when temperatures rose above the

estimated frost point. Gates et al. (2002) argue that, for

this to occur, the water surface layer of the particles is

either coated with HNO3 or combined with HNO3 to form

nitric acid trihydrate (NAT). In either case, the HNO3

would serve to inhibit evaporation of the volatile aerosol

component. Popp et al. (2002) measured ~30 parts per

billion (ppb) HNO3 in the Athena II wake, likely sufficient to explain the observed persistence of condensed

volatiles. It is not clear if HNO3 is produced during SRM

combustion, during intense chemical processing in the

plume wake, or by a combination of the two processes.

In any case, the Athena II observations indicate that SRM

plume wake particles are more complex and likely to be

more reactive than previously considered. 

Space launch activity has undergone substantial

growth over the last decade (60% increase per decade;

FAA, 2000). If this growth rate is realized through SRM

use and is sustained over the next several decades, the

associated ozone depletion from stratospheric chlorine

and particle injection might become significant. Under

some plausible scenarios, ozone loss due to rockets may

become greater than that due to chlorofluorocarbons by

the year 2050. However, recent technology developments

indicate that growth in hydrocarbon engine emissions may

exceed growth in SRM emissions. The change in calculated ozone impact resulting from SRM replacement by

hydrocarbon engines cannot be confidently estimated at

this time.

4.8.4 Future Tropospheric Ozone

The future evolution of tropospheric ozone was

considered in some detail in the recent Third Assessment

Report (TAR) of the Intergovernmental Panel on Climate

Change (IPCC, 2001), and here we focus primarily on the

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results presented therein. Research groups involved in

3-D tropospheric chemistry modeling were invited to participate in the TAR through a model intercomparison

activity. Ten different models submitted results for both

the year 2000 (Y2000), and for one scenario for the year

2100 (Y2100). The groups were asked to run with a standard set of emissions for NOx, CO, and volatile organic

compounds (VOCs) for 2000 and for abundances of longlived gases; the standard simulation for 2000 provided the

baseline from which changes in greenhouse gases were

calculated.

The model simulations for 2000 were tested by

comparison with the seasonal cycle of ozone at five locations (20°N-75°N) for 700, 500, and 300 hPa, and the simulation for CO was compared with surface concentrations

from five sites (40°S-80°N). The models captured the

broad seasonal patterns for CO and ozone, but there was a

wide dispersion among the models for both gases. Most

models were within 30% of the ozone observations.

However, it should be noted that there were large variations in the calculation of the present-day tropospheric

ozone budget. In particular, there is disagreement between

models on the sign of the net photochemical production,

possibly driven by differences in the calculated flux of

ozone from the stratosphere.

A large number of scenarios for future emissions

were developed for the IPCC Assessment, but only one of

these, designated the Special Report on Emissions

Scenarios (SRES) A2p, was used for the Y2100 standard

calculation; a CH4 abundance of 4.3 parts per million by

volume (ppmv) was specified for Y2100. The TAR notes

that the SRES scenarios assume few controls on CO, NOx,

and VOCs, and that A2p has overall the highest emissions.

It was chosen so that once the response of ozone and OH

to these extreme emissions was understood, other scenarios and intermediate years could be interpolated with

some confidence.

The increase in the tropospheric column of ozone

from Y2000 to Y2100 calculated by the ten 3-D models

was 16.6 to 26.5 DU, and the TAR adopted a best estimate

of 22 DU. Although the increases in the tropospheric

ozone column given by the various models were rather

consistent, the spatial and vertical distributions of the

ozone changes varied considerably from model to model.

A footnote added in proof to the TAR notes that these “tropospheric” column changes erroneously included the

lower stratosphere, and that the troposphere-only changes

are 25-33% less, or about 16 DU. The present-day column

of tropospheric ozone is about 34 DU, so the calculated

increase is a substantial fraction of the current amount in

the troposphere, and it is a few percent of the total column

of midlatitude ozone.

The TAR used the standard Y2000 and Y2100 calculations, along with three other cases (Y2000 with a 10%

increase in CH4, Y2100 with no increase in NOx emissions,

and Y2100 with no increase in NOx, VOC, and CH4) to

define simple linear relationships for the absolute change

in tropospheric ozone and the relative change in OH as a

function of the CH4 abundance, and the emission rates for

NOx, CO, and VOC. These relationships were then used

to estimate the evolution of six different emission scenarios

from 2000 to 2100. The changes in the tropospheric ozone

column for 2000 to 2100 for the six scenarios are from –4

to +22 DU, whereas the changes for 2000 to 2050 are from

4 to 15 DU. Clearly, the scenario selected for the Y2100

model study represents an extreme case.

The calculations that were carried out for the TAR

did not allow for changes in stratospheric ozone, nor did

they allow for changes in climate. Calculations with the

coupled ocean-atmosphere-chemistry model at the Hadley

Centre (U.K.) have shown that the increases predicted for

tropospheric ozone and CH4 by 2100 are smaller in simulations that allow for climate change than in those that do

not (Johnson et al., 2001). Trace gas emissions were taken

from the same scenario, A2, as in the TAR, but CH4 was

calculated based on emissions. In the control run, ozone

at 650 hPa increased from 40 ppbv to 65 ppbv at midlatitudes, whereas in the run with changing climate, ozone

increased to only 49 ppbv. The increase in CH4 was from

1.67 ppmv in 1990 to 3.65 ppmv in 2100 in the control

run, but to only 3.23 ppmv in the climate change run. Most

of the difference for both gases occurred after 2040. The

main causes for smaller changes in ozone and CH4 in the

climate change run were (1) an increase in photochemical

destruction of ozone caused by higher water vapor, and

(2) a decrease in the CH4 lifetime caused by the higher

temperatures increasing the rate coefficient for OH + CH4,

and by higher OH.

While these results indicate that estimates of future

tropospheric ozone may be reduced in a future climate,

there are other feedbacks that may result in higher tropospheric ozone levels. For example, increased deep convection in a future climate could lead to increases in NOx

production by lightning (e.g., Toumi et al., 1996). Climate

change may also lead to increases in temperaturedependent emissions (e.g., isoprene) or changes in the flux

of ozone from the stratosphere, resulting in additional

increases in tropospheric ozone. Current uncertainties in

our knowledge about such feedbacks hinder our ability to

predict future changes.

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4.9 SYNTHESIS OF CURRENT

UNDERSTANDING OF PAST AND 

FUTURE CHANGES IN OZONE

Table 4-5 summarizes the important processes that

we believe have contributed to the observed past changes

in global ozone, our estimated level of scientific understanding of these processes, and their likely importance in

causing the observed trend in column ozone. Our level of

scientific understanding is given on a scale from low to

high. The impact on global ozone is rated from negligible

to important. The table also summarizes the processes

that will affect the future evolution of the ozone layer, an

estimate of the likely importance of that process, and a

measure of our confidence in that prediction. The stated

confidence ranges from high to low. The importance for

column ozone ranges from minor to important.

4.9.1 Discussion of Past Changes

There is very high confidence that gas-phase chlorine chemistry is the main driver of ozone trends in the

upper stratosphere. This is based on overall reasonable

agreement between observed and modeled ozone trends

in the upper stratosphere, in terms of vertical profile

(Figure 4-30) and latitudinal structure (e.g., WMO, 1999,

Figure 6-20). Although there are substantial inter-model

differences in the magnitude of the trends near 40-45 km

for the simulations examined here (Figure 4-30), these are

strongly related to differences in modeled temperature

trends in the upper stratosphere, and also to changes in

constituents such as CH4 (although a quantitative analysis

of these effects is beyond the scope of this Assessment).

In the upper stratosphere, lower temperatures cause a slow

down of the gas-phase chemical loss reactions due to chlorine, while CH4 controls the ClO/HCl partitioning (as is

apparent in observations in Figure 4-29), and therefore

larger CH4 trends lead to smaller ozone trends. Note that

long-term global observations of upper stratospheric CH4

over 1979-2000 are unavailable to compare and constrain

the model results. Upper stratospheric CH4 measurements

from HALOE during 1991-2000 show substantial interannual variability and changes in “trends” that are not well

understood at present (Randel et al., 1999), and these are

difficult to extrapolate to the longer 1979-2000 period.

Although the processes responsible for upper stratospheric

ozone losses are reasonably well understood, the contribution to column ozone trends is relatively small: only

15-30% of the extratropical column ozone losses for 1980-

2000 originate at altitudes above 25 km.

In the lower stratosphere, there is fairly high confidence that halogen-related chemistry is responsible for a

large fraction of the midlatitude ozone trends. This is

based on the fact that 2-D and 3-D models with imposed

halogen trends are able to approximately simulate the latitudinal, vertical, and seasonal characteristics of observed

midlatitude ozone trends. There are important caveats to

this attribution, because the 2-D models all overestimate

observed ozone trends in the tropics and underestimate

the magnitude and seasonality of trends in the midlatitude

NH, and there is relatively large inter-model variability

for trend calculations in the midlatitude SH. The underestimate of NH midlatitude trends during winter-spring

may be due to a number of factors: effects of water vapor

trends and uncertainties in bromine amount and chemical

reaction rates, lack of detailed polar vortex chemistry and

mixing to midlatitudes in 2-D models, and the contribution of dynamically induced ozone changes in observations. Also, the precise mechanisms for the decadal-scale

ozone changes that occur in the models are not simply isolated, and the relative importance of polar processing (and

export to midlatitudes) versus in situ activation awaits

clarification. The overestimates of tropical ozone trends

may be related to relatively weak dynamical isolation of

the tropics in many models (Hall et al., 1999; see also

Figure 4-28), and the large inter-model variability in SH

extratropics is associated with simulation of the Antarctic

ozone hole and transport out of the vortex.

Although the models are able to approximately

simulate the 20-year trends in midlatitudes, both 2-D and

3-D simulations are less successful at simulating the actual

time series of midlatitude ozone. This is particularly the

case in the SH, where models overestimate SH ozone

losses after the eruption of Mt. Pinatubo and miss substantial ozone declines observed during the middle 1980s.

Variability in the wave-driven Brewer-Dobson circulation

contributes to interannual ozone changes in the lower

stratosphere during winter and spring, and this factor is

relatively more important in the NH where wave driving

is stronger and highly variable. Although the contribution of this process to decadal ozone trends is difficult to

quantify, it probably accounts for 30% or less of the

observed NH trends during 1979-2000.

The lowermost stratosphere is the region where the

ozone trends, and their causes, are least certain. Long

records of ozonesonde data are limited to relatively few

stations over NH midlatitudes. The role of chemistry is

difficult to quantify because of low temperatures and

occasionally high humidity. This makes the role of heterogeneous chemistry on aerosols and cirrus particles

uncertain. Dynamics plays a relatively large role for

ozone variability in the lowermost stratosphere. This

region is sensitive to the Brewer-Dobson circulation (from

above), and empirical studies over NH midlatitudes show

strong relationships between ozone and changes in tropoGLOBAL OZONE

4.70
GLOBAL OZONE

4.71

Table 4-5. Summary of understanding of past long-term midlatitude stratospheric ozone changes and expected importance in the future.

Influence Process Level of Scientific Region of Local Contribution to Past Likely Importance for

Understanding Change Column Changes and Future Column Change

(LOSU) a Sign of Forcing b Over Next 50 Years

ChemicalCFCs, halons, Cly/Bry processing from vortex Medium Tropopause-30 km Important; negative Decreasing over time

and other Cl/Br Cly T-dependent gas-phase Very high 25-50 km Important; negative Decreasing over time

source gases chemistry in US

Cly/Bry gas-phase chemistry High Tropopause-25 km Important; negative Decreasing over time

in LS/LMS

N O2 5 + H O on sulfate aerosol High Tropopause-30 km Important; negative Decreasing over time 2

Cly processing on cold liquid aerosol Medium Tropopause-20 km Important; negative Decreasing over time

CH4 CH4 chemistry High Tropopause-50 km Minor; positive Minor

HOx chemistry High Tropopause-20 km Minor; negative Minor

and 30-50 km

Effect on H O and Medium Tropopause-20 km Minor; negative Minor 2

heterogeneous chemistry

N O NO 2 y chemistry Very high 25-45 km Minor; negative Increasing over time

H O HO 2 x and heterogeneous chemistry Medium Tropopause-50 km Minor; negative UnknownAircraft NOx, HOx, aerosols Medium Tropopause-20 km Minor; positive Minor

Rockets Cly, alumina, soot Medium Tropopause-50 km Negligible; negative Minor

Dynamical and radiativeStratospheric planetary Transport Medium 15-30 km Important; positive, Important

waves negative

Tropospheric circulation Stratosphere-troposphere Medium-low Tropopause-20 km Important; positive, Important

exchange and LMS dynamics c negative

Radiatively active species Thermal structure Medium-high Tropopause-50 km Minor; positive (US), Important

(CO2, CH4, N O, H 2 O) 2 negative (LS/LMS)

GHG direct effect d

OtherSolar irradiance Modulation of photolysis rates Medium Tropopause-50 km Negligible; positive, Minor

negative

Volcanoes Heterogeneous chemistry Medium Tropopause-25 km Important; negative Transient; possibly largeTropospheric O3 precursors RO2-NOx chemistry Medium Surface-tropopause Minor; positive Importanta The meaning of LOSU for dynamical influences is different from that used for the chemical influences. The relevant physics andmechanisms of the stratospheric planetary wave forcings are well understood. However,

it is hard to quantify their past effect in the real atmosphere. For tropospheric circulation changes, uncertainties come from our lack of clear understanding of the mechanisms controlling tropopause height. For directGHG changes, uncertainties come from factors influencing how well the changes in H O near the tropopause are known. For future changes, the LOSU for all processes is limited by the quality of predictive models 2 . b The meaning of “forcing” for the effects of stratospheric planetary waves and tropospheric circulation on ozone is different than for chemical forcings. Here “forcing” is used in the sense that decadal changes in these

dynamical properties of the atmosphere have occurred that can cause decadal changes in ozone. The possible causes of these dynamical changes are discussed in Section 4.6.

c Tropopause height is a proxy for LMS dynamics. d GHG direct effect means the radiative impact of changes in GHGs and any consequent changes in the thermal structure of the atmosphere GHG effect. The direct greenhouse effect of GHGs is a forcing in the same way as

the chemical forcings are defined, namely, an increase in atmospheric concentrations of certain gases as a result of anthropogenic activity has an effect on the atmosphere that can, in principle, be separated and determined.
spheric circulation and tropopause height (forcing from

below). The interrelationships between ozone and circulation parameters in the lowermost stratosphere are poorly

quantified at present, and estimates of circulation effects

on decadal changes in ozone are highly uncertain.

Overall, the balance of evidence still suggests that

increases in halogens have been the most important driver

for long-term changes in global ozone. However, compared with previous Assessments, the longer observation

time series now available have emphasized other factors,

such as dynamical variability, which modulate ozone in

the lower part of the stratosphere.

4.9.2 Discussion of Future Changes

Over the next 50 years the stratospheric burden of

chlorine is expected to return to the same level as that

which existed in 1980 (Chapter 1). Accordingly, the reversal of the halogen forcing of the observed past trends

in global ozone is expected to lead to an ozone increase.

However, a number of other factors may affect the rate

and extent of this recovery from the halogen-induced

depletion.

In the upper stratosphere, where past trends have

been due to chlorine chemistry, ozone is expected to return

to 1980 values well before 2050. Coupled 2-D model calculations suggest that a cooling of the stratosphere (due

to GHG increases) will significantly increase US ozone

through a slowdown of the gas-phase chemical loss reactions (Figure 4-43). Without this cooling, similar models

predict that US ozone would not have returned to 1980

levels by 2050, highlighting the important impact of this

feedback.

In the lower stratosphere, the decrease in stratospheric chlorine will also lead to a recovery from the

observed depletion of the past two decades. Although

bromine has also contributed to LS O3 changes, the reduction in chlorine is expected to dominate future trends. A

range of 2-D models indicate that NH and SH midlatitude

column ozone (which is dominated by the LS) will return

to 1980 levels between 2025 and 2050, although there is a

large uncertainty in this timing (Figure 4-42). The models

that include the effect of stratospheric cooling on gasphase chemistry indicate that this will also significantly

accelerate the rate of column ozone increase in this period.

An important caveat to these predictions is that the 2-D

models examined here do not include a detailed treatment of polar processes and mixing to midlatitudes.

Heterogeneous processing of chlorine species in the polar

vortex may significantly affect midlatitude ozone, and

over the next decade or so stratospheric cooling could

enhance polar ozone loss (see Chapter 3). Therefore, in

the near future continued midlatitude ozone losses may

still be observed.

Midlatitude ozone will be most susceptible to

enhanced chemical loss by volcanic eruptions in the next

20 years or so while stratospheric chlorine levels remain

high. During this period, a large eruption, similar in magnitude to that of Mt. Pinatubo in 1991, may cause a similar transient decrease in column ozone that lasts a few

years. Predicted changes in N2O and CH4 will have a relatively small impact on column ozone over the next 50

years.

Any future changes in stratospheric circulation and

transport could have the potential to affect global column

ozone. At present it is not possible to predict what these

changes may be and therefore what effect they may have

on column ozone. This is a major limitation on our predictive capability.

Beyond 2050 the stratosphere will return to one in

which halogens play less of a role in controlling the distribution of ozone. Ozone trends may then be dominated

by trends in other gases that produce chemical species

known to catalytically destroy ozone (e.g., H2O, N2O, and

CH4). For example, under some possible scenarios, future

increases in N2O may lead to enhanced loss due to

enhanced NOy chemistry. However, it seems that the

radiative cooling of future increased GHGs, which will

lead to increased US ozone, may dominate. In the future

low-halogen atmosphere, the cooling of the stratosphere

will also have less leverage to destroy ozone through polar

processing. Combined with the fact that model studies

predict an increase in tropospheric ozone by 2100, it seems

likely that by the end of this century the atmosphere may

contain more ozone than in the recent past.

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Appendix 4A

DESCRIPTION OF OZONE DATASETS

A brief description of the various observational ozone datasets used in this Assessment is included here. More

detailed discussions can be found in the Report of the International Ozone Trends Panel: 1988 (WMO, 1990) and previous Assessments.

4A.1 Ground-Based Total Ozone Measurements 

Three types of ground-based instruments are used for long-term monitoring of total ozone content. Their characteristics and performance have been discussed in numerous WMO Assessments (e.g., WMO, 1995, 1999), and therefore

only a brief description is provided here. The longest records of continuous reliable measurements are available from

stations equipped with Dobson spectrophotometers. The first regular Dobson measurements started in the 1920s, and a

worldwide network of Dobson instruments was established in the 1950s and early 1960s. Recent intercomparisons show

that at present Dobson instruments at main sites agree with the World Standard Dobson 83 within 0.5% (WMO, 1995,

1999). 

The Brewer spectrophotometer, developed in the early 1980s, can measure total ozone with approximately the

same uncertainty as the Dobson spectrophotometer. All Brewer instruments are calibrated against the traveling standard

Brewer instrument, and the traveling standard is calibrated against the Brewer triad at Toronto (Kerr et al., 1997). The

Dobson and Brewer calibrations are independent, and the fact that there is no systematic difference between the two

types of instruments adds to confidence in the data quality. 

Filter ozonometers are widely used in the former Soviet Union countries, and long-term reliable records are available from 1972. This instrument is less accurate than the Dobson and Brewer instruments, and the calibration is traceable

to the Dobson reference. Filter ozonometer observations are important for estimations of long-term ozone variations

over vast regions where no Dobson and Brewer data are available and where observations from several stations can be

averaged to reduce the random errors (Bojkov et al., 1994). 

Total ozone data available from the World Ozone and Ultraviolet Radiation Data Center (WOUDC) were used

here, including stations with reliable data and without any major calibration problems in their records. Fioletov et al.

(1999) presented a recent review of the ground-based network performance. Data are available from 45 stations with

continuous records starting before 1983 (WMO, 1999), and a larger set (up to 109 stations) was used to estimate zonal

and global ozone averages for the period 1979-2001, following the methodology described in Bojkov and Fioletov

(1995).

4A.2 Satellite Total Ozone Measurements and Merged Datasets 

TOTAL OZONE MAPPING SPECTROMETER (TOMS)

In a TOMS measurement, ozone is derived using the “backscattered albedo,” the ratio of the radiance backscattered from the Earth-atmosphere system to the extraterrestrial solar irradiance. The solar irradiance is measured using an

onboard diffuser plate. Instrument change should cancel when using this ratio, except for change in the diffuser plate

itself, something that proved to be a problem for Nimbus-7 (N7). This was solved by using a triple diffuser for the

Meteor-3 (M3) and Earth Probe (EP) instruments. All the TOMS datasets used for this report have been processed using

the version 7 (v7) algorithm. Each TOMS instrument has had unique hardware problems, but in-orbit calibration techniques have been used to maintain the relative calibration of each instrument to approximately 1%/decade accuracy. For

long-term trend determination there is the additional uncertainty of the relative calibration of EP TOMS, M3 TOMS, and

N7 TOMS.
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NIMBUS-7 TOMS: NOVEMBER 1978 TO APRIL 1993 

The absolute accuracy of the Nimbus-7 (N7) TOMS instrument depends on the radiometric calibration that was

performed pre-launch, using standard lamps and laboratory diffuser plates. The time-dependent (in-orbit) calibration is

maintained by a technique in which the ratio of two long-wavelength channels (non-ozone sensitive) is stabilized.

Although instrument changes in N7 TOMS were becoming significant by the end of its life, the calibration was maintained to within ~1%/decade. This is supported by a comparison with a network of 39 Dobson stations that have nearly

complete coverage over the entire N7, M3, EP period (McPeters and Labow, 1996). This comparison showed that the

drift of N7 TOMS relative to this Dobson network was less than 0.5%/decade, with an average offset of 0.83%. 

It is strongly suspected that there is an error in the N7 TOMS data due to “toggling.” In approximately 1983, N7

TOMS began to exhibit small errors in the synchronization of the chopper wheel and the electronics. The instrument

appeared to “toggle” between two possible states, and when the calibration was set, one of the two states had to be chosen

as the “correct” state. It now appears that the incorrect choice was made, resulting in an error of somewhat less than 1%

between 1982 and 1988. This error has little effect on the overall trends for this instrument, but because of the accident

of its timing it is likely that it increases the apparent amplitude of any solar cycle term in a regression model. This error

will be corrected when these data are reprocessed with a v8 algorithm.

METEOR-3 TOMS: AUGUST 1991 TO DECEMBER 1994 

The absolute calibration of the Meteor-3 (M3) TOMS instrument was explicitly set to be the same as that of the N7

TOMS instrument. During the period of overlap when both instruments were operating (8/91-4/93) a normalization

factor was derived for the M3 instrument based on a subset of the data for which strict match-up criteria in space and time

were met. This makes the M3 dataset a consistent extension of the N7 dataset rather than a totally independent dataset. 

The time dependence of M3 TOMS was independently maintained. This was the first instrument to have a triple

diffuser assembly: a cover diffuser, a working diffuser that was used weekly, and a reference diffuser that was exposed

only a few times a year. This approach was very successful; the well-protected reference diffuser showed no detectable

signs of degradation after years of operation. (The same is true of the Earth Probe diffuser after more than 5 years in

space.) Precession of the orbit of the Meteor-3 spacecraft led to periods of lower quality data when the orbit was near the

terminator, but the effect of these missing data on long-term changes should be minimal since these periods are distributed evenly through the data record.

EARTH PROBE TOMS: JULY 1996 TO DECEMBER 2000

The pre-launch radiometric calibration procedure used for the Earth Probe (EP) TOMS instrument was described

in detail by McPeters et al. (1998). Comparison of EP TOMS ozone with that from the 39 Dobson stations through April

2000 shows an average difference of 1.5%. However, a wavelength calibration performed on the QuikTOMS instrument

infers that there were small wavelength errors in the EP instrument, which will reduce the EP TOMS ozone values by

about 0.5% when a planned reprocessing is done. This will bring the difference with the Dobson measurement down to

about 1%, and make the EP data consistent with N7 to within about 0.2%, well within the expected calibration errors of

the two instruments.

As with the M3 instrument, EP maintains its time-dependent calibration using a triple diffuser plate. The calibration

appears to have been very stable through 1999, but in 2000 artifacts began to appear in the TOMS data related to the look

angle. By the summer of 2000, ozone when measured to the left of the orbital track was on average 6 DU lower than when

measured to the right of the orbital track. Analysis showed that this cross-track bias was caused by real changes in the

optical properties of the scan mirror. The principle that instrument errors should cancel in the radiance-to-irradiance ratio

works to some extent, but the angle of the scan mirror when it is viewing the diffuser plate is quite different from when it is

looking at the Earth, leading to residual error. It is expected that this scan angle-dependent error will be corrected when the

data are next reprocessed. The EP TOMS data should be fully usable for analysis of long-term change through the end of

1999, and an empirical scan bias correction has been derived that should make the EP data usable through the end of the

year 2000. However, by 2001 the changes in the optical properties of the scan mirror have become so large that the EP

data should not be used for trend determination at this time.
SOLAR BACKSCATTER ULTRAVIOLET (SBUV, SBUV/2) 

The long-term collection of total ozone measurements from the Solar Backscatter Ultraviolet (SBUV) spectrometer began with its launch on the NASA Nimbus-7 (N7) spacecraft in 1978. With the demonstrated success of this instrument, the National Oceanic and Atmospheric Administration (NOAA) adopted it (with some modifications) for launch on

the NOAA afternoon operational satellite series beginning in 1985, and designated it SBUV/2. Thus far, the NOAA

series has included NOAA-9, -11 and –14; NOAA-16 was launched in late 2000. The SBUV retrieval algorithm provides

both total ozone and the vertical distribution of ozone in ~5-km-thick layers, and the data here use the version 6 backscatter

ultraviolet (BUV) retrieval discussed in Chapter 1 of the SPARC/IOC report (SPARC, 1998). The differences in total

ozone between instruments are approximately 1 to 2%. The temporal overlap of instruments can be used to account for

calibration biases between instruments, and Miller et al. (2002) have derived a continuous, intercalibrated dataset incorporating measurements from Nimbus-7 (11/78 to 2/85), NOAA-9 (3/85 to 12/88 and 1/94 to 12/95), NOAA-11 (1/89 to

12/93 and 7/98 to 12/00), and NOAA-14 (1/96 to 6/98). The separate time periods for NOAA-9 and -11 data result from

the precession of the satellite orbits, and the corresponding loss of BUV measurements as the orbit drifts from local day to

night.

There can be substantial biases in the SBUV/2 retrieved ozone for measurements made at very high solar zenith

angle (SZA), and the gridded analyses used here only included data for SZA up to 80°. For the “normal” afternoon orbit

of SBUV/2, this results in a loss of data at high southern latitudes during local winter (data are unavailable poleward of

~55°S). However, the combination of this SZA cutoff and the orbital drifts of NOAA-9 and -11 results in increased data

loss at high southern latitudes for some years, namely, 1988, 1992, 1993, and 1998, when winter data are unavailable

poleward of ~45°S. This undersampling by the SBUV+SBUV/2 data is one potential source of difference with the other

datasets over SH midlatitudes.

MERGED TOMS+SBUV/2 

The Goddard merged ozone dataset was derived using measurements from six satellite instruments: the Nimbus-7

TOMS and SBUV; the NOAA-9, -11, -14, and -16 SBUV/2; and the Earth Probe TOMS. The D-pair (306 nm-311 nm)

measurements from SBUV and SBUV/2 instruments were used to stabilize their calibrations as a function of time. The

overlaps of these instruments with N7 and EP TOMS were used to establish the relative calibration of each instrument.

These were then all adjusted to the first 2 years of EP TOMS as a reference. During this process it was discovered that the

N7 TOMS exhibited time-dependent calibration shifts, in part due to a 1.5-DU error in 1984 when the instrument went

into the “toggling” mode. The N7 TOMS calibration has been adjusted for these calibration shifts, leading to a deduced

trend over its lifetime that is smaller than that shown in previous studies. The data, and information about how they were

constructed, can be found at http://code916.gsfc.nasa.gov/Data_services/merged.

NIWA ASSIMILATED DATASET

The National Institute of Water and Atmospheric Research (NIWA) assimilated total ozone dataset is derived from

combining measurements obtained by TOMS instruments flown on the Nimbus-7, Meteor-3, Earth Probe, and ADEOS

(Advanced Earth Observing Satellite) satellites, together with the Global Ozone Monitoring Experiment (GOME) instrument flown on the European Remote Sensing (ERS)-2 satellite (Burrows et al., 1999). A description of these datasets and

their homogenization is provided in Bodeker et al. (2001). Briefly, comparisons between TOMS data and measurements

from the ground-based Dobson spectrophotometer network were used to remove offsets and drifts in the Nimbus-7 and

Earth Probe satellite data. Statistical model fits to the Nimbus-7 and Meteor-3 differences were used to adjust the Meteor3 TOMS data. Adjusted Earth Probe TOMS data were similarly used to correct ADEOS TOMS data and GOME data. 

COMPARISON OF COLUMN OZONE DATASETS

Some systematic differences between the separate total ozone datasets described above were found. Figure 4A-1

shows the 1979-1987 annual mean percentage differences between the zonal averages estimated from four satellite

datasets and the ground-based data as a function of latitude. All datasets agree to within 1% of the ground-based data at

40°-50°N. However, larger differences of 2-4% are seen in the tropics and the Southern Hemisphere for the TOMS,

SBUV-SBUV/2 ,and merged satellite datasets (the NIWA dataset uses adjustments to fit the ground-based data, and hence

GLOBAL OZONE

4.87
the mean differences are small). Satellite data are 5-6% higher than ground-based measurements over the Antarctic. For

the entire 60°S-60°N zone the area-weighted differences between the other datasets and ground-based zonal averages are

the following: 1.5% higher for the TOMS, 2.1% higher for the merged satellite dataset, 0% for the NIWA assimilated

dataset, and 3% higher for the SBUV-SBUV/2 dataset. 

Deseasonalization removes systematic differences between the datasets, and the deseasonalized total ozone deviations agree to within a root mean square (rms) difference of 0.6-0.8 % (depending on the dataset) over the 1979-2001

period. However, the differences vary in time, as shown in Figure 4A-2, where the individual data records are compared

with the average of all five datasets. The range of differences is typically ~1%, but can be as high as 2% at certain times

(notably 1991-1993). The SBUV-SBUV/2 dataset is on the low end of the ensemble for much of the period after 1989.

4A.3 Ozone Profile Measurements

Long-term ozone profile datasets include satellite, ground-based Umkehr, and ozonesonde (balloon) measurements. The satellite data are near-global, whereas the Umkehr and ozonesonde measurements are taken at only a few stations, mostly located in northern midlatitudes. The primary satellite data source is derived from the Stratospheric Aerosol

and Gas Experiment (SAGE I and SAGE II) measurements, which cover the altitude region ~15-50 km. Although ozone

profile information is also available from SBUV-SBUV/2 satellite measurements, there are important uncertainties in the

long-term calibration and changes associated with orbital precession of the NOAA-9 and -11 spacecraft, and therefore a

homogeneous long-term profile dataset useful for trend analyses is not available at present. Detailed descriptions and

intercomparisons of the SAGE, Umkehr, and ozonesonde measurement techniques and instruments can be found in

SPARC (1998), and only brief summaries are provided here.

GLOBAL OZONE

4.88

-60 -40 -20 0 20 40 60

Latitude

- 4

- 2

0

2

4

6Differ ence (%)

TOMS zonal means

Merged satellite data

SBUV-SBUV/2

NIWA assimilated data 1978 1980 1982 1984 1986 1988 1990 1992 1994 1996 1998 2000 2002

-2

-1

0Deviation (%)

G round-based dataset

T OMS zonal means

SBUV-SBUV/2

Merged satellite data

NIWA assimilated data set

6 0oS - 60oN

Figure 4A-1. The difference between total ozone zonal

means estimated from four datasets and ground-based

data, as a function of latitude for the period 1979-1987.

Monthly means calculated from the ground-based data

have been subtracted from the monthly means of each

of the other four datasets, and annual averages calculated; results expressed in percent of the ground-based

means. From Fioletov et al. (2002).

Figure 4A-2. Differences between the five estimates

of deseasonalized ozone anomalies over 60°N-60°S

(shown in Figure 4-2) and their average. The average

of all five estimates was calculated for each season,

and then subtracted from all five datasets. Updated

from Fioletov et al. (2002). 
SAGE I 

The SAGE measurement technique (McCormick et al., 1989) is based on solar occultation, with ozone profile

measurements obtained at sunrise and sunset on each of 14 orbits per day. This technique provides high vertical resolution (~1 km) and very small long-term drifts resulting from instrument calibration. However, spatial sampling is limited,

and it takes approximately one month to sample the latitude range 60°N to 60°S. SAGE I ozone profile data cover the

time period February 1979 to November 1981. A difficulty in combining the SAGE I and SAGE II data for trend studies

arises from an apparent error in the reference altitude for SAGE I. An empirical altitude correction has been applied to

the SAGE I data in an attempt to remove this bias (H.J. Wang et al., 1996).

SAGE II 

SAGE II ozone profile data commenced in November 1984 and continue to the present (Decmeber 2002). The

analysis here used SAGE II version 6.1 (v6.1), which was released in March 2002. It includes a new algorithm for altitude registration, which results in a large reduction in the scatter in the transmission profile. There is evidence that the

vertical resolution has been significantly improved over previous versions and is now approximately 0.75 km; v6.1 is

placed on a 0.5-km altitude grid. The species separation and inversion in v6.1 is similar to that used in v5.96 (the basis

for the SAGE II ozone data in WMO (1999), but v6.1 incorporates an aerosol model that improves the ozone retrievals in

the presence of enhanced aerosol (H.J. Wang et al., 2002). The primary change to v6.1 was the method of estimating the

species errors, which now include terms for the Rayleigh scattering and aerosol model uncertainties. In general, the

ozone results are of good quality down to the tropopause. The data analyzed here exclude time periods of enhanced

aerosol extinction following the eruption of Mt. Pinatubo in June 1991.

UMKEHR OZONE PROFILE MEASUREMENTS

Ground-based measurements use the Umkehr technique for measuring the profile of ozone concentration. It consists of observing a series of ultraviolet wavelengths of solar radiation, some of which are strongly absorbed by atmospheric ozone. A sequence of radiance measurements is made as the Sun rises or sets, and information about the ozone

profile is obtained from an inversion algorithm applied to the measurements (e.g., Mateer and Deluisi, 1992). A new

Umkehr retrieval algorithm (ALG-99) has been developed by Petropavlovskikh et al. (2001), with results discussed

extensively in Bojkov et al. (2002). ALG-99 utilizes latitudinally and seasonally dependent first-guess ozone profiles.

Unlike previous versions, the new first-guess profiles do not depend on total ozone, and therefore the influence of the

total ozone trend on the retrieval has been largely reduced.

Relatively long records of Umkehr measurements (exceeding 15-20 years) are available from 13 stations. However,

time series from many stations have significant discontinuities resulting from occasional changes in instrumentation.

Careful analysis of the individual time series in addition to station histories suggests that the long-term data are of sufficient quality at only a relatively few stations. The results here are based on data from Arosa, Switzerland (47°N); Belsk,

Poland (52°N); Boulder, Colorado, U.S. (40°N); and Haute-Provence, France (44°N).

OZONESONDES

A network of stations provides ozone profile information from balloonborne sondes, which measure ozone from

the ground to about 30 km, with a vertical resolution of ~150 m. Details of ozonesonde measurements and data processing are discussed in SPARC (1998). One important detail is that each sonde profile is normalized to coincident total

ozone measurements, and the amount of data available for trend analyses depends on the chosen limits of this normalization factor (SPARC, 1998; Logan et al., 1999). Data are available to the end of 2000 for all the stations used in the

SPARC (1998) analysis, which include 14 locations spanning the latitude range 32°N-75°N. The Uccle dataset has been

reprocessed since the SPARC study. The ozonesondes at Uccle switched from Brewer Mast to electrochemical concentration cell (ECC) sondes in March 1997, with a transition period during which both sondes were used, to allow adjustment for any discontinuity in the dataset.

GLOBAL OZONE

4.89
GLOBAL OZONE

Appendix 4B

2-D MODEL SCENARIOS

Table 4B-1. Definition of the 2-D model scenarios. All model runs used the halocarbon scenario described

in Table 4B-2 below. IPCC refers to IPCC (2001).

Scenario Period Motivation CH4 N2O CO2 Aerosol

MC2 1979-2000 Best past trend IPCC IPCC IPCC D. Considine

MA2 2000-2050 Control A2 IPCC A2 IPCC A2 IPCC A2 Background

MB2 2000-2050 Control B2 IPCC B2 IPCC B2 IPCC B2 Background

MLCH4 2000-2050 Low CH4 0.8 ¥ A2 IPCC A2 IPCC A2 Background

increase post2000

MHCH4 2000-2050 High CH4 1.2 ¥ A2 IPCC A2 IPCC A2 Background

MLN2O 2000-2050 Low N2O IPCC A2 0.8 ¥ A2 IPCC A2 Background

MHN2O 2000-2050 High N2O IPCC B2 1.2 ¥ A2 IPCC B2 Background

MLCO2 2000-2050 Low CO2 IPCC A2 IPCC A2 0.8 ¥ A2 Background

MHCO2 2000-2050 High CO2 IPCC A2 IPCC A2 1.2 ¥ A2 Background

MEPI 2000-2050 Episodic aerosol IPCC A2 IPCC A2 IPCC A2 Episodic

MENH 2000-2050 Enhanced back- IPCC A2 IPCC A2 IPCC A2 Enhanced

ground aerosol

Table 4B-2. Halocarbon scenario used in the 2-D model calculations. This scenario is essentially the same

(except for a few small deviations) as the baseline scenario Ab described in Tables 1-13 and 1-16 of Chapter 1.

Time CFC- CFC- CFC- CFC- CFC- CCl4 CH3CCl3 HCFC- HCFC- HCFC- Halon Halon Halon Halon CH3Br CH3Cl

11 12 113 114 115 22 141b 142b 1211 1202 1301 2402

1950 0.39 5.81 0.19 1.13 0.00 34.54 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 7.02 483.42

1955 2.96 14.59 0.53 2.39 0.00 41.16 0.05 0.25 0.00 0.00 0.00 0.00 0.00 0.00 7.09 492.78

1960 9.16 29.81 1.18 3.63 0.00 51.13 1.44 1.28 0.00 0.00 0.00 0.00 0.00 0.00 7.26 510.97

1965 23.26 59.14 2.44 4.90 0.00 63.63 4.59 3.82 0.00 0.10 0.00 0.00 0.00 0.00 7.44 528.77

1970 52.62 114.74 4.93 6.37 0.12 75.36 15.96 10.11 0.00 0.21 0.02 0.00 0.00 0.01 7.66 540.55

1975 105.99 203.64 9.84 8.14 0.50 85.04 39.20 21.17 0.00 0.30 0.12 0.00 0.04 0.04 7.93 546.48

1980 162.26 295.85 19.52 10.28 1.33 93.03 78.53 36.72 0.00 0.36 0.42 0.01 0.24 0.11 8.20 549.05

1985 205.27 382.18 36.72 12.64 2.91 99.62 105.68 56.32 0.00 0.51 1.04 0.02 0.74 0.20 8.55 550.10

1990 256.34 474.07 67.97 15.36 5.06 105.67 124.91 81.33 0.00 1.24 2.27 0.03 1.66 0.31 8.95 550.51

1995 270.62 524.80 83.23 16.18 7.49 103.66 108.63 111.78 2.09 5.27 3.25 0.04 2.33 0.42 9.40 550.68

2000 262.64 540.48 81.70 16.38 8.97 97.94 49.81 140.15 10.81 10.61 3.97 0.05 2.82 0.41 9.31 550.72

2005 250.66 541.09 78.01 16.41 9.15 92.25 21.85 172.56 23.71 18.91 4.14 0.02 3.17 0.37 8.81 550.72

2010 236.07 522.50 73.87 16.27 9.15 80.13 10.79 180.55 34.65 24.28 3.76 0.01 3.40 0.31 8.44 550.72

2015 218.15 497.02 69.68 16.07 9.13 66.11 5.67 168.54 39.14 26.12 3.19 0.00 3.52 0.26 8.44 550.72

2020 198.88 472.78 65.70 15.84 9.10 54.55 2.75 141.32 36.78 24.67 2.59 0.00 3.56 0.21 8.17 550.72

2025 179.90 449.72 61.95 15.59 9.07 45.01 1.22 110.25 30.68 21.49 2.05 0.00 3.52 0.17 8.17 550.72

2030 162.01 427.79 58.41 15.34 9.04 37.13 0.51 84.70 24.33 18.19 1.59 0.00 3.44 0.13 8.17 550.72

2035 145.52 406.92 55.07 15.09 9.01 30.64 0.21 65.05 18.85 15.20 1.22 0.00 3.33 0.10 8.17 550.72

2040 130.51 387.08 51.93 14.84 8.98 25.28 0.08 50.08 14.38 12.56 0.93 0.00 3.19 0.08 8.17 550.72

2045 116.94 368.20 48.96 14.60 8.95 20.85 0.03 38.03 10.80 10.26 0.70 0.00 3.04 0.06 8.17 550.72

2050 104.72 350.24 46.16 14.36 8.92 17.21 0.01 26.56 7.55 8.06 0.52 0.00 2.88 0.05 8.17 550.72

4.90
Table 4B-3. GHG scenarios (from IPCC, 2001). Scenarios A2 and B2 were used in the model runs. CO2 is

from the Integrated Science Assessment Model (ISAM) reference CO2 scenario (p. 807 of IPCC, 2001).

Time CH4 (ppbv) N2O (ppbv) CO2 (ppmv)

IS92a A2 B1 B2 IS92a A2 B1 B2 IS92a A2 B1 B2

1970 1420 1420 1420 1420 295 295 295 295 326 325 325 325

1980 1570 1570 1570 1570 301 301 301 301 338 337 337 337

1990 1700 1700 1700 1700 308 308 308 308 354 353 353 353

2000 1810 1760 1760 1760 319 316 316 316 372 369 369 369

2010 1964 1861 1827 1839 328 325 324 323 393 390 388 388

2020 2145 1997 1891 1936 339 335 333 328 418 417 412 408

2030 2343 2163 1927 2058 350 347 341 333 446 451 437 429

2040 2561 2357 1919 2201 361 360 349 338 476 490 463 453

2050 2793 2562 1881 2362 371 373 357 342 509 532 488 478

GLOBAL OZONE

4.91
CHAPTER 5

Surface Ultraviolet Radiation: Past and Future

Lead Authors:

J.B. Kerr

G. Seckmeyer

Coauthors:

A.F. Bais

G. Bernhard

M. Blumthaler

S.B. Diaz

N. Krotkov

D. Lubin

R.L. McKenzie

A.A. Sabziparvar

J. Verdebout

Contributors:

A. Arola

V.E. Fioletov

J.R. Herman

P.W. Kiedron

J.B. Liley

C.S. Long

S. Madronich

T.J. Martin

R. Meerkötter

P.J. Neale

H. Slaper

J.R. Slusser

P. Taalas

M. van Weele

E.C. Weatherhead

C.S. Zerefos

CHAPTER 5

SURFACE ULTRAVIOLET RADIATION: PAST AND FUTURE

Contents

SCIENTIFIC SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.1

5.1 INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.5

5.2 CURRENT UNDERSTANDING OF FACTORS AFFECTING SURFACE UV RADIATION . . . . . . . . . . . . 5.5

5.2.1 Dependence of UV on Geophysical Variables . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.5

5.2.1.1 The Extraterrestrial Solar Spectrum . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.6

5.2.1.2 Dependence of UV on Ozone and Other Trace Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.6

5.2.1.3 Dependence of UV on Clouds . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.6

5.2.1.4 Dependence of UV on Aerosol Scattering and Absorption . . . . . . . . . . . . . . . . . . . . . . . . . . 5.7

5.2.1.5 Dependence of UV on Surface Albedo . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.8

5.2.1.6 Dependence of UV on Altitude . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.8

5.2.1.7 Angular Dependence of UV . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.9

5.2.1.8 Dependence of UV Under Water . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.9

5.2.2 Daily Forecasting of Ozone and UV . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.10

5.2.3 Remaining Questions and Uncertainties . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.11

5.3 AVAILABLE RESOURCES FOR STUDYING SURFACE UV RADIATION . . . . . . . . . . . . . . . . . . . . . . . . 5.11

5.3.1 Ground-Based Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.11

5.3.1.1 Spectroradiometers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.11

5.3.1.2 Broadband Filter Radiometers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.12

5.3.1.3 Narrowband Multifilter Radiometers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.13

5.3.1.4 Data Quality and Databases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.13

5.3.1.5 Instrument Intercomparisons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.14

5.3.2 Radiative Transfer Models and Model Validation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.15

5.3.3 Satellite Estimates and Validation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.16

5.3.4 Statistical Modeling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.19

5.4 UV CLIMATOLOGY, TEMPORAL CHANGES, AND TRENDS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.20

5.4.1 Ground-Based UV Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.20

5.4.2 Estimates from Satellite Data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.22

5.4.3 Estimates from Radiative Transfer Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.23

5.4.4 Estimates from Ground-Based Ancillary Data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.23

5.5 EXPECTATIONS OF UV IN THE FUTURE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.25

5.5.1 Links with the Recovery of the Ozone Layer . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.25

5.5.2 Factors Related to Climate Change . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.28

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.31

APPENDIX 5A: SPECTRAL DATA AVAILABLE FROM DATABASES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.43

5A.1 World Ozone and Ultraviolet Radiation Data Centre (WOUDC) . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.43

5A.2 European Ultraviolet Database (EUVDB) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.44

APPENDIX 5B: INTERNET ADDRESSES FOR UV SITES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.45

5B.1 General UV Information . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.45

5B.2 International UV Projects . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.45

5B.3 Radiative Transfer Codes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.45
5B.4 Extraterrestrial Spectra . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.45

5B.5 General UV Index Information . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.46

5B.6 Internet Sites with UV Information by Country . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.46
SCIENTIFIC SUMMARY

Important Confirmations of Previous Results

• Additional measurements have confirmed that decreases in column ozone lead to increases in surface ultraviolet

(UV) radiation. The relationship has been further documented from existing and new sites. Spectral data records

from 65 stations are now available in international data centers. Episodes with elevated UV irradiance associated

with low total ozone continue to occur.

• There is clear evidence that the long-term UV changes are not driven by ozone alone, but also by changes in cloudiness, aerosols, and surface albedo. The relative importance of these factors depends on local conditions.

• Enhanced values of UV radiation continue to be observed at high latitudes in the Southern Hemisphere under the

Antarctic ozone hole. Highest biologically weighted UV doses under the ozone hole are typically not observed in

October when maximum ozone depletion occurs, but in November and early December when solar elevations are

higher and low ozone values still prevail. Changes in the duration and spatial extent of the Antarctic ozone hole are

therefore more important for Antarctic UV levels than the annual ozone minimum.

New Findings

• Satellite estimates of surface UV radiation from the Total Ozone Mapping Spectrometer (TOMS) dataset have been

compared with ground-based measurements at several more sites since the previous Assessment. In general the

estimates capture short-term and long-term variability. However, the estimates are systematically higher than

ground-based measurements at many sites. The differences in monthly average erythemal (“sunburning”) UV irradiance range from about 0% at some clean sites and up to 40% in the Northern Hemisphere. The fact that the agreement is better at the cleaner sites suggests that the differences are due to aerosols and/or pollutants near the ground.

• New algorithms have been developed that produce UV maps with improved spatial (down to 1 km) and/or temporal

resolution (down to 30 minutes). These algorithms take into account the effects of clouds, surface albedo, topographical features, and aerosols by using information from other satellite and ground datasets, in addition to satellite column ozone data. The detailed maps have contributed to a better understanding of the geographical and temporal variability in surface UV radiation and have enabled specific UV impact studies. The achievable agreement

between the results of such algorithms and ground-based measurements for monthly averages is within ±5%.

• The length of ground-based spectral UV measurement has reached more than 10 years for some sites. It has been

shown that a record of 10 years is still too short to derive statistically significant trends.

• UV increases associated with the ozone decline have been observed by spectral measurements at a number of sites

located in Europe, North America, South America, Antarctica, and New Zealand.

• Calculations based on pyranometer (total irradiance), total ozone, and other meteorological measurements have

been used to reconstruct surface UV irradiance at several mid- to high-latitude sites. Pyranometer and other meteorological data serve as proxies for parameters affecting UV other than ozone. The reconstructed datasets, which

extend backward in time to as early as the 1960s at some sites, show long-term increases in erythemal irradiance of

about 6-14% over the last 20 years. At some sites approximately half of the changes can be attributed to total ozone

changes. It is believed that the increases of UV irradiance derived from the ground-based reconstructed data are

clear indicators of the long-term changes that have occurred since the 1980s. However, these reconstructions contain several assumptions on the nature of radiative transfer, are not measurements of UV irradiance, and are not representative on a global scale.

• The reconstructed datasets show that the number of hours of UV levels above certain threshold values has increased

at some northern midlatitude sites since the 1970s. For example, in Toronto the number of hours per year with the

UV Index above 7 has doubled.

SURFACE ULTRAVIOLET RADIATION

5.1
• Snow cover can increase surface UV irradiance by more than 50%. The presence of snow several kilometers away

from the observing site may still lead to significant enhancement. Three-dimensional radiative transfer models

have shown that the region of significance (defined by increases of UV irradiance by more than 5%) can extend

beyond a radius of 40 km.

• In the Antarctic, ozone depletion has been the dominant factor for increases in UV irradiance. The future evolution

of UV radiation is therefore expected to follow the ozone recovery. However, because of changes in other influencing factors, such as changes in cloud cover, aerosols, or snow/ice cover, UV radiation may not return exactly to

pre-ozone-hole values.

• Elsewhere, including the Arctic, the impact of other influencing factors can be comparable to the impact of ozone

depletion. The large uncertainties in future changes of these prevent reliable predictions on the future evolution of

UV irradiance. Furthermore, climate-change-induced trends in cloudiness, aerosols, and snow/ice cover are

expected to be seasonally and geographically dependent, leading to differences in future UV irradiance in different

parts of the world.

• The 2001 Intergovernmental Panel on Climate Change report (IPCC, 2001) states that decreases in sea ice cover

over the Northern Hemisphere have occurred and are likely to continue. Furthermore, increases in global cloud

cover are likely. Such changes lead to decreases in UV irradiance over large areas of the globe. Quantitative estimations of these effects and their regional dependence cannot yet be made. Although UV irradiance above the

surface is enhanced by snow-covered areas, a decrease in sea ice and snow cover will result in an increase of UV

dose for organisms living under water and on land areas previously covered by snow.

Advances in Our Understanding

• The effects of climate change on surface UV irradiance are twofold. The first effect is indirect and results from climate changes that influence total ozone. The second effect is direct and results from changes in other climatic

variables such as clouds, aerosols, and snow cover.

• Spectral irradiance and actinic flux calculated by several radiative transfer models agree to within 2% when input

assumptions are properly chosen. For cases when the Sun is close to the horizon and for shorter wavelengths, the

relative standard deviations are somewhat higher (5%), but then irradiance is very small.

• The influence of tropospheric aerosols on UV irradiance may be larger than previously thought, and may affect

large areas of the globe. This result is based on studies on the reduction of UV irradiance by aerosols using groundbased instruments and aircraft missions.

• Three-dimensional radiative transfer models have improved the understanding of atmospheric scattering and

absorption processes. These models are used to study the effects of inhomogeneous features such as broken and

scattered clouds, and nonuniform surface albedo.

• Advances have been made in the understanding of the penetration of UV radiation under water that have allowed

the estimation of the underwater UV environment on a global scale using combined satellite datasets.

• A reanalysis of TOMS satellite data with respect to the influence of changes in cloudiness over Europe has confirmed the result found previously by ground-based measurements showing that pyranometric irradiance at the

ground has decreased in recent decades. Thus it can be concluded that UV increases due to ozone are partly masked

by the increased cloudiness in some regions.

• Uncertainties of ground-based UV measurements are now found to be greater than those estimated in the previous

Assessment. A rigorous uncertainty analysis for spectroradiometric solar UV measurements has been performed.

For a typical well-maintained instrument it has been shown that uncertainties are as high as ±12.7% at 300 nm and

SURFACE ULTRAVIOLET RADIATION

5.2
±6.3% for erythemally weighted irradiance. Major reasons are uncertainties related to calibration standards and

wavelength uncertainties. Deviations of about 5-10% between well-maintained instruments and between measurements and models have been seen at recent intercomparisons. In general, broadband instruments have even higher

uncertainties than spectroradiometers.

• Better calibration methods have been developed that result in smaller deviations among spectroradiometers. New

instruments have been developed and existing instruments have improved with regard to stability, wavelength

alignment, and angular response.

• The extraterrestrial solar UV spectrum can be retrieved by Langley plot analysis of ground-based measurements at

very clean sites. New studies have shown that these retrievals agree with space-based measurements to within 3%.

SURFACE ULTRAVIOLET RADIATION

5.3

SURFACE ULTRAVIOLET RADIATION

5.5

5.1 INTRODUCTION

A major concern regarding a decrease in stratospheric ozone is the consequential increase of solar ultraviolet (UV) radiation passing through the atmosphere and

reaching the Earth’s surface. Ozone absorbs radiation

strongly in the UV, and the presence of ozone and oxygen

in the stratosphere results in the absorption of nearly all

solar radiation below 290 nm. Thus virtually all UV-C

radiation (200-280 nm) does not reach the troposphere or

the Earth’s surface. Solar UV-B radiation (280-315 nm)

is significantly absorbed by atmospheric ozone, whereas

only a small fraction (less than 3%) of UV-A radiation

(315-400 nm) is absorbed by ozone. Accurate measurement of spectral UV has historically been a difficult task.

The fact that irradiance varies by many orders of magnitude over a relatively short wavelength range (290-320

nm) requires that useful instruments have a wide dynamic

range and a high degree of spectral purity and accuracy.

Also the long-term stability of UV instruments and their

absolute calibration standard are still difficult to maintain. Consequently, good quality routine spectral measurements did not start until the late 1980s, and these longer

records are few in number.

Solar UV radiation that reaches the ground is influenced by many complicated scattering and absorption

processes in the atmosphere and at the Earth’s surface.

These processes must be understood in order to describe

fully the characteristics of the current climatology of surface UV radiation on a global scale and to project it to the

past and future. Section 5.2 presents findings since the

previous Assessment (WMO, 1999) that have improved

the understanding of radiative transfer (RT) processes

defining the magnitude, wavelength dependence, and

angular distribution of UV radiation at and above the

Earth’s surface and under water or ice. Section 5.2 also

discusses the UV Index, a direct application of our understanding of the processes defining surface UV. Section

5.3 discusses resources that are applied to understand and

quantify RT processes. These resources include groundbased measurements of surface UV radiation and other

pertinent variables; satellite-based measurements of extraterrestrial solar flux, ozone, cloud cover, and reflectivity;

RT computer models; and statistical models that quantify

dependencies of UV radiation on scattering and absorption. Assessment of the effectiveness and accuracy of

these resources (also Section 5.3) includes intercomparisons of different ground-based instruments, intercomparison of RT models, and comparisons of the surface measurements with satellite estimates and RT model results.

In addition to an assessment of recent findings

described in the scientific literature, this chapter provides

the link between ozone depletion as assessed in earlier

chapters of this report and the impacts of increases of surface UV radiation that are assessed in the most recent

United Nations Environment Programme (UNEP)

“Effects Panel” report (UNEP, 2002). In general, aquatic

and terrestrial biological systems (including human

beings), material degradation, and pollution photochemistry are sensitive to UV radiation, and in most of the cases

the sensitivity increases with decreasing wavelength. The

wavelength dependence of the sensitivity of a particular

biological system to UV radiation is defined by an action

spectrum, and the erythemal (sunburning) action spectrum as standardized by the Commission Internationale

de l’Éclairage (CIE; McKinlay and Diffey, 1987) is often

used. Studies on impacts of UV radiation require knowledge of the present UV climatology (including average

and extreme values) and any changes that have occurred

in the past. Results of studies showing that there have

been past changes in UV (caused by changes in ozone as

well as changes in other variables such as clouds, aerosols,

and snow cover) are given in Section 5.4. Section 5.5

presents estimates of future surface UV radiation determined from predictions of ozone changes given in earlier

chapters. Effects of other climatological changes, as

determined and discussed in the recent Intergovernmental

Panel on Climate Change (IPCC) report (IPCC, 2001),

that may directly influence surface UV radiation in the

future are also discussed in Section 5.5.

5.2 CURRENT UNDERSTANDING OF FACTORS

AFFECTING SURFACE UV RADIATION

Since the previous Assessment (WMO, 1999),

available resources have been applied to advance our

understanding of processes that affect surface UV radiation and have recently been reviewed (e.g., Taalas et al.,

2000a; Blumthaler and Taalas, 2001). New research

results regarding the dependencies of UV on geophysical

variables are presented in Section 5.2.1. Following that is

a discussion on the UV Index forecast (Section 5.2.2),

which is a direct application of our knowledge and

resources. Validation of these daily forecasts allows an

ongoing opportunity to assess the accuracy of our

resources. Finally, Section 5.2.3 briefly summarizes

aspects regarding surface UV radiation that we do not

fully understand.

5.2.1 Dependence of UV on Geophysical

Variables

One of the main objectives for making good quality

measurements of surface UV radiation is to understand

the dependencies of UV on absorption and scattering

processes that occur in the atmosphere and at the Earth’s
SURFACE ULTRAVIOLET RADIATION

5.6

surface. In principle, if the spatial distribution of all

absorbers and scatterers within the atmosphere and at the

Earth’s surface were fully known, it would be possible to

determine the wavelength dependence and angular distribution of surface UV radiation by model calculations. In

practice, the complicated distribution of the predominant

variables (clouds, ozone, aerosols, and surface albedo)

and their interactive impact on UV irradiance makes

detailed calculations on a global scale an extremely difficult task.

5.2.1.1 THE EXTRATERRESTRIAL SOLAR SPECTRUM

Accurate knowledge of the solar spectrum is important for studying surface UV, since it is used in RT models

that are ultimately compared with measurements. At the

time of the previous Assessment (WMO, 1999) several

satellite measurements were available showing agreement to within ±3% (Cebula et al., 1996). Appendix 5B

(Section 5B.4) lists Internet sites that provide information

on measured extraterrestrial spectra.

Knowledge of the stability of solar irradiance at

UV wavelengths is important because it drives many geophysical processes, including the formation of stratospheric ozone. It could also be used as a reference standard for UV measurements. Changes in solar irradiance

from 1700 to present were determined by Fligge and

Solanki (2000). Their model assumes that solar irradiance variations on time scales from days to centuries are

due only to the changing distribution of solar surface magnetic features. The study determined an increase of solar

spectral irradiance at solar activity cycle minimum since

the Maunder minimum (year 1700) of 3.0% for wavelengths less than 300 nm and 1.3% for the band 300-400

nm. The total (all wavelengths) and the visible irradiance

showed an increase of about 0.3%. Rozema et al. (2002)

note that a past long-term increase in solar UV-C radiation would lead to an increase in the production of stratospheric ozone that would reduce surface UV-B irradiance.

5.2.1.2 DEPENDENCE OF UV ON OZONE AND OTHER

TRACE GASES

The effects of the absorption of atmospheric ozone

on surface UV radiation were well understood at the time

of the previous Assessment (WMO, 1999), and there has

been no major change in our knowledge since then. The

wavelength dependence of the sensitivity in UV irradiance to ozone had been fairly well established both by

observations and comparison of observations with RT

models (e.g., Fioletov et al., 1997; Bodhaine et al., 1998),

and these relationships continue to be confirmed (e.g.,

Bartlett and Webb, 2000; Casale et al., 2000; Cho et al.,

2000; Miyauchi et al., 2000). Further work has been done

regarding the effects of variations of the vertical distribution and the temperature of ozone (Sabziparvar et al.,

1998; Krzyscin, 2000; Lapeta et al., 2000). These effects

could influence surface UV should there be a long-term

change in the vertical profile or effective column temperature of ozone; however, the effects are relatively small

compared to past or expected future long-term changes in

total ozone.

The effects of sulfur dioxide on surface UV were

shown to be negligible on a global scale but could be significant at sites near local pollution sources or continuous

volcanic activity (Fioletov et al., 1998), or over wider

areas in the aftermath of major volcanic eruptions. Other

gases that absorb in the UV (e.g., nitrogen dioxide, nitric

acid, and formaldehyde) are not significant under natural

conditions, but could be significant under heavy pollution

situations.

5.2.1.3 DEPENDENCE OF UV ON CLOUDS

Clouds have more influence on surface UV irradiance than any other atmospheric variable. Although

important, cloud effects on UV irradiance are difficult to

quantify. The effect of clouds on UV is understood in

principle. However, in practice the necessary parameters

used to calculate local cloud effects are rarely available,

and if they were, the complexities of cloud geometry need

to be specified in sufficient detail and require the use of

three-dimensional (3-D) model calculations.

Under overcast conditions, clouds decrease the irradiance measured at the surface (Josefsson and Landelius,

2000; Renaud et al., 2000). However, enhancements of

up to 25% can occur under broken cloud conditions

(Estupiñán et al., 1996; Sabburg and Wong, 2000; Weihs

et al., 2000), or if there are reflections from cloud decks

below high-altitude observation sites such as Mauna Loa

Observatory (McKenzie et al., 2001b). Even for large

cloud fractions, the reduction in irradiance can be small if

the clouds do not obscure the direct beam. Thus, one of

the most important parameters is whether or not the Sun

is obscured (Grant and Heisler, 2000; Schwander et al.,

2002). For individual sites, this poses difficulties for satellite products. When a histogram of cloud transmission is

plotted as a function of cloud amount, a bimodal distribution typically results (e.g., Seckmeyer et al., 1997;

McKenzie et al., 1998b; McKenzie et al., 2001b), with a

lower peak resulting from conditions when cloud obscures

the Sun, and a higher peak corresponding to conditions

where clouds do not block the Sun. Since the Sun can be

unobscured even for large cloud fractions, or obscured

even for small cloud fractions, the quantification of cloud
SURFACE ULTRAVIOLET RADIATION

5.7

effects can become problematic (e.g., Udelhofen et al.,

1999; Matthijsen et al., 2000). There are also complications when the scattering of radiation by clouds enhances

effects such as absorption by ozone (e.g., Fioletov et al.,

1997, 2002) or scattering by aerosols (e.g., Erlick et al.,

1998; Mayer et al., 1998b) within the cloud.

The presence of scattered or broken clouds poses

difficulties for comparisons between ground-based measurements and satellite estimates of surface UV irradiance.

In this situation direct solar radiation is either obscured or

not obscured by a cloud at the ground-based measurement

site, whereas the satellite measures an average cloud

amount over its footprint.

Ground-based measurements of cloud cover are

now available at several sites using automated all-sky

imagery (Sabburg and Wong, 2000). Although in some

cases high thin clouds may not be distinguishable in

all-sky images, these continuous records of the spatial

distribution of clouds used in parallel with UV irradiance

measurements offer the potential to understand and quantify cloud effects more accurately.

5.2.1.4 DEPENDENCE OF UV ON AEROSOL

SCATTERING AND ABSORPTION

Aerosols are highly variable over space and time.

They attenuate UV flux through the atmosphere to an

extent that is mostly described by the aerosol optical depth

(AOD) and the average column value of the single scattering albedo (wo), which is the ratio of scattering to

extinction (where extinction = scattering + absorption).

The wavelength dependence of the AOD is generally

assumed to be proportional to l-a, where l is wavelength,

and a is the Angstrom coefficient. Measurements of AOD

at visible and UV-A wavelengths using sunphotometry

are routinely carried out (e.g., Schmid et al., 1997, 1999).

Currently the worldwide Aerosol Robotics Network

(Holben et al., 1998) is the best source of information on

aerosol particle size, optical depth, and single scattering

albedo. Significant progress has been made recently in

the measurement and quantification of the effects of

aerosols at UV-B wavelengths.

One of the objectives of the Photochemical Activity

and Solar Ultraviolet Radiation (PAUR) campaign held in

Greece during June 1996 and the PAUR II campaign in

Greece and Italy during May-June 1999 was to study the

optical properties of aerosols. The AOD in the UV was

shown to be approximately inversely proportion to wavelength (i.e., a = 1.0 ± 0.5), and values of wo between 0.84

and 0.98 were determined by comparing model results

with the measurements (Marenco et al., 1997; Kazantzidis

et al., 2000; Kylling et al., 1998). A variety of atmospheric gaseous and particulate concentrations were measured with radiative and optical characteristics of the

atmosphere during the PAUR II campaign (Zerefos et al.,

2002). Kouvarakis et al. (2002) found that the AOD and

single scattering albedo correlated well with surface measurements of ammonium sulfate, indicating the key role

that this gas plays in radiative forcing in the area.

Variations of calcium originating from the Sahara also

correlated well with AOD measurements. The campaigns

also studied how changes in stratospheric ozone might

influence tropospheric photochemistry through changes

in the UV environment (Jonson et al., 2000; Zerefos et al.,

2001; Balis et al., 2002; Hofzumahaus et al., 2002; Zanis

et al., 2002). These impacts are discussed in the Effects

Panel report (UNEP, 2002).

Wenny et al. (2001) studied the variations in AOD

at 317, 325, 332, and 368 nm from July through December

1999 using a multifilter radiometer near Asheville, North

Carolina, and found the mean optical depth was 0.33 at

368 nm and increased to 0.40 at 317 nm, indicating a value

of about 1 for a. The study used an RT model to investigate the reduction in erythemal UV irradiance relative to

a baseline case with AOD = 0.2 and wo = 1.0. When the

AOD was increased to 1.3, maximum reduction in UV

was 44% for wo = 0.75, 29% for wo = 0.90, and only 18%

for wo = 1.0. Vermeulen et al. (2000) have developed a

method for retrieving the scattering and microphysical

properties of atmospheric aerosols (including wo) from

measurements of solar transmission, aureole, and angular

distribution of the scattered and polarized sky light in the

solar principal plane.

It has been shown by ground-based measurements

that large reductions of UV-B occur under absorbing

aerosols such as smoke from biomass burning (e.g., Ilyas

et al., 2001; Kirchhoff et al., 2001), forest fires (e.g.,

McArthur et al., 1999), or desert dust (di Sarra et al.,

2002a). Recent work suggests that anthropogenic aerosols

that absorb in the UV region may play a more important

role in attenuating UV irradiances than has been assumed

previously (Jacobson, 2001).

Comparisons between satellite-derived UV and

ground-based spectral measurements have revealed inconsistencies in satellite-derived UV that are probably related

to the inability of the satellite sensors to correct for

boundary layer extinctions. Ground-based estimates of

regional UV irradiances suffer from a similar inability to

correct for horizontal inhomogeneities in boundary layer

extinction as well as inhomogeneities in the troposphere

and stratosphere. Only at the cleaner sites is there good

agreement within the experimental errors. At continental

sites in the Northern Hemisphere, the satellite-derived UV
SURFACE ULTRAVIOLET RADIATION

5.8

estimations are too large (McKenzie et al., 2001a; Fioletov

et al., 2002).

Gröbner et al. (2001) made a comparison of AODs

determined by seven Brewer spectrophotometers and a

Li-Cor spectrophotometer at Huelva, Spain, between 1

and 10 September 1999. For all instruments, measurements of AOD at 306, 310, 313.5, 316.7, and 320 nm

agreed to within ±0.03 over a range of AOD between 0.2

and 0.6 at 306 nm. These results suggest that the existing

Brewer network can be used to determine UV AODs, an

important consideration for extending the coverage of

ground-based measurements for studying scattering/

absorption processes and long-term trends. Routine measurements of AOD in the UV using Brewer instruments

have been reported (e.g., Meleti and Cappellani, 2000;

Kerr, 2002).

5.2.1.5 DEPENDENCE OF UV ON SURFACE ALBEDO

The presence of snow cover surrounding an

observing site increases UV irradiances (McKenzie et al.,

1998a; Minschwaner, 1999; Herman and McKenzie et al.,

1999; Krotkov et al., 2001; Chubarova et al., 2002;

Fioletov et al., 2002), even when the snow is several kilometers away (Degünther et al., 1998; Degünther and

Meerkötter, 2000b; Weihs et al., 2001). Recently developed RT models have shown that certain snow distributions can cause significant (defined as being >5%)

increases of surface UV irradiance even if the snow is farther than 40 km from a site (Degünther et al., 1998;

Ricchiazzi and Gautier, 1998; Smolskaia et al., 1999;

Lenoble, 2000; Mayer and Degünther, 2000). Models

have also been used to study the complicated interactions

between snow enhancements combined with atmospheric

scattering and absorbing processes (Aoki et al., 1999;

Renaud et al., 2000; Krotkov et al., 2001).

Snow enhancements between 8% and 39% for surface UV at 324 nm were reported in the previous

Assessment (Herman and McKenzie et al., 1999).

Increases are greater at Arctic sites surrounded by uniform snow cover and less at urban sites or sites near open

water (Fioletov et al., 2002). Site-specific enhancements

are attributable to increases in the “regional” albedo,

which represents a spatial average around the site. The

regional albedo may be determined from the snow

enhancement values (Schwander et al., 1999; Krotkov et

al., 2001; Chubarova et al., 2002) and is a function of the

age and depth of snow as well as the terrain and other features around the site. Measurements of UV irradiance

(320-325 nm) made at Syowa, Antarctica (Aoki et al.,

2000), were compared with those made at four Japanese

sites (Sapporo, Tsukuba, Kagoshima, and Naha) for the

same values of total ozone and solar zenith angle (SZA).

Values at Syowa were observed to be about 60% higher

than the upper limit of values seen over Japan (Takao et

al., 1999), making the albedo enhancement an important

consideration when combined with low total ozone values

seen under the Antarctic ozone hole.

Information regarding the average albedo in the

satellite field of view is important for space-based estimates of surface UV (Krotkov et al., 2001). When snow

depth information is used as input to estimate regional

albedo, better agreement between ground-based measurements and satellite estimates is achieved, particularly

under cloudy conditions (Arola et al., 2002; Krotkov et

al., 2002). However, day-to-day knowledge of regional

albedo on a global scale remains a challenge.

Combining measurements with RT model results

can be used to derive estimates for unknown input parameters by statistical fitting procedures. When aerosol

optical depth is relatively low (and thus the effect of inaccurately known optical characteristics of aerosols is

small), regional albedo of partly snow-covered terrain can

be derived (Schwander et al., 1999; Gröbner et al., 2000;

Weihs et al., 2001). Values of regional surface albedo in

the range of 0.3 to 0.9 were determined, depending on

local conditions of topography and snow coverage.

5.2.1.6 DEPENDENCE OF UV ON ALTITUDE

Since the previous Assessment (WMO, 1999), the

dependence of UV irradiance on altitude has been further

quantified. In practice the altitude dependence of UV irradiance is itself dependent on differences in surface albedo,

boundary layer extinctions by aerosols, and tropospheric

ozone concentrations. Therefore the dependence is not

represented by a single value (Seckmeyer et al., 1997).

Even when these effects are ignored, it has been found

that at higher altitude the dependence of UV irradiance on

SZA and wavelength changes. In general, for erythemally

weighted UV, irradiances in clean conditions increase

between 5 and 10% per kilometer (McKenzie et al.,

2001b; Zaratti et al., 2002; Schmucki and Philipona,

2002), with the greatest increase occurring at SZA ~ 60-

70°. To model the dependence correctly at Mauna Loa

Observatory, it was necessary to consider the effects of

sky irradiance scattered from below the observatory

(McKenzie et al., 2001b). In mountainous regions the

vertical gradient in some instances can be larger (up to

50% per kilometer) because of local effects such as

increasing albedo with altitude and high concentrations

of ozone or aerosols in the lower troposphere (Seckmeyer

et al., 1997). Larger altitude gradients in the free tropo-
sphere have also been measured from aircraft under polluted situations over Greece (Varotsos et al., 2001).

5.2.1.7 ANGULAR DEPENDENCE OF UV

Ground-based measurements of UV irradiance are

generally made with a horizontal diffuse surface that

would ideally follow a response proportional to the cosine

of the angle from normal (vertical) incidence. In many

cases the collection efficiency of the diffuser falls below

the cosine function for large SZA, and calculations are

made to adjust (usually increase) the measured irradiance

(Bais et al., 1998a; Chubarova and Nezval, 2000; Fioletov

et al., 2002). The adjustments depend on the angular distribution of incident radiation and are complex functions

of SZA, wavelength, total ozone, ozone distribution,

aerosol, and cloud amounts. The relationships of measured to adjusted values are corrected by use of RT models.

For many biological and photochemical processes,

actinic fluxes rather than cosine-weighted irradiances

are more appropriate. However, such measurements have

not been generally available until quite recently

(Hofzumahaus et al., 1999, 2002; Shetter and Müller,

1999). Recent progress has been made in converting irradiances to actinic fluxes, offering the prospect of deriving

historical changes of actinic flux from the extensive

existing database of irradiance measurements (Kazadzis

et al., 2000; McKenzie et al., 2002; Webb et al., 2002). It

was found that the most important parameters defining

the relation between actinic flux density and global UV

irradiance are the ratio of direct to global irradiance,

whether or not the Sun is obscured, and the description of

the angular distribution of the diffuse radiation. Under

clear skies, retrieved spectral actinic flux densities agree

with model calculations to within ±5% for a variety of

aerosol conditions. The assumption of an isotropic distribution of the diffuse radiation can lead to a wavelengthdependent overestimation of actinic flux densities from

10 to 15%, depending on SZA, especially in atmospheres

with high aerosol content (Kazadzis et al., 2000).

A campaign to measure spectral global UV irradiance and actinic flux at the ground, supported by ancillary measurements used to characterize the atmosphere,

was carried out in August 2000 at Nea Michaniona, in

northern Greece (Webb et al., 2002). It was established

that the ratio of actinic fluxes (F) to horizontal irradiance

(E) is between 1.4 and 2.6 for UV wavelengths. This ratio

is a function of wavelength, SZA, and the optical properties of the atmosphere. Both the wavelength and SZA

dependency of the ratio decrease when the scattering in

the atmosphere increases and the direct beam proportion

of global irradiance decreases, as expected.

5.2.1.8 DEPENDENCE OF UV UNDER WATER

The underwater UV environment is an important

consideration for studies of the sensitivities of aquatic

(both freshwater and saltwater) species to UV radiation

(de Mora et al., 2000). These effects are discussed extensively in the UNEP Effects Panel report (UNEP, 2002).

There have been recent spectral measurements of UV

under fresh water (e.g., Sommaruga and Psenner, 1997;

Laurion et al., 1997; Bukaveckas and Robbins-Forbes,

2000; Markager and Vincent, 2000) and salt water (e.g.,

Booth and Morrow, 1997; Kuhn et al., 1999). Measurements of UV penetration into seawater using a

biochemical deoxyribonucleic acid (DNA) dosimeter

combined with a spectroradiometer have been reported

(Boelen et al., 1999). In general, these measurements

show that there is wavelength-dependent absorption by

water in the UV that increases with decreasing wavelength. The underwater absorption has strong dependence on the abundance of dissolved organic matter that

has wide temporal and spatial variability. A sensitivity

study conducted by RT modeling has shown that the main

parameters controlling levels of the most harmful UV-B

radiation underwater for clear-sky conditions are the SZA,

seawater bio-optical properties, and total ozone amount

(Vasilkov and Krotkov, 1997). Attenuation of UV-B irradiance and DNA dose rate with water depth is primarily

controlled by the total seawater absorption coefficient and

its spectral dependence.

Estimates of UV radiation penetration into the

ocean waters are now available on a global scale by combining Total Ozone Mapping Spectrometer (TOMS) satellite estimates of UV irradiance at the ocean surface with

the SeaWiFS (Sea-viewing Wide Field-of-view Sensor)

satellite ocean-color data and using a model to define seawater optical UV properties (Vasilkov et al., 2001).

Weekly maps of underwater UV irradiance and DNAweighted exposure are calculated using monthly-mean

SeaWiFS chlorophyll and diffuse attenuation coefficient

products, daily SeaWiFS cloud fraction data, and daily

maps of TOMS-derived surface UV irradiance. The final

products include global maps of weekly-average UV-B

irradiance and DNA-weighted daily exposures at depths

of 3 m and 10 m, and depths where the UV-B irradiance

and DNA-weighted dose rate at local noon are equal to

10% of their surface values.

Global mapping of underwater UV radiation creates many new challenges. The challenges are mostly

related to larger uncertainties in physical input parameters caused by biological processes within the oceans.

Vasilkov et al. (2002) discuss the problems encountered

in the assessment of the underwater UV irradiance from

SURFACE ULTRAVIOLET RADIATION

5.9
space-based measurements, and propose approaches to

overcome the difficulties by combining different satellite

datasets (TOMS, SeaWiFS, and Moderate Resolution

Imaging Spectroradiometer (MODIS)). Both SeaWiFS

and MODIS provide some estimate of seawater optical

properties in the visible. Currently, the problem of accurate extrapolation of visible data to the UV spectral range

is not solved completely, and there are only a few available measurements. Vasilkov et al. (2002) propose to

parameterize UV absorption by oceanographic constituents empirically by using bio-optical measurements

from a variety of ocean waters. Another problem is the

lack of reliable data on pure seawater absorption in the

UV. Laboratory absorption measurements of the middle

UV for both pure water and pure seawater are required

(Fry, 2000).

The transmission of UV radiation through sea ice

is reduced by absorption and scattering processes that

occur within and on top of the ice. UV radiation is strongly

absorbed by both colored dissolved organic matter and

particulate organic matter, including ice algae. Scattering

occurs above the ice by snowcover, within the ice, and

from interstitial brine (Vasilkov et al., 1999). Consequently, the transmission changes during the season,

either declining by an order of magnitude as a highly scattering turbid brine layer develops on top of the ice, and as

ice algal communities bloom (Trodahl and Buckley, 1990;

Perovich et al., 1998), or increasing if snow melt ponds

form on top of the ice, thereby reducing albedo and

scattering (Belzile et al., 2000). In the latter case, the

under-ice spectrum becomes enriched in UV-B relative to

photosynthetically active radiation, and may exacerbate

UV-Beffects.

5.2.2 Daily Forecasting of Ozone and UV

Forecast of the UV Index is a direct application of

the resources used to study UV radiation and the knowledge of the dependencies of surface UV on absorption and

scattering processes. Forecast values of total ozone are

used as input to RT models or statistical models. Most of

the agencies that currently produce forecasts of the UV

Index had been doing so by the time of the previous

Assessment (WMO, 1999) in accordance with the international definition and standardization (WMO, 1994,

1997). Although many agencies provide UV Index forecasts on their web pages, only a few have published their

methodology (Burrows et al., 1994; Long et al., 1996;

Burrows, 1997; Bais et al., 1998b; Lemus-Deschamps et

al., 1999). Many countries have adopted and/or modified

the above methodologies to produce their own UV Index

forecasts. Further information regarding the UV Index

can be found on Internet sites listed in Appendix 5B

(Section 5B.5).

With the increase of computing speeds, more countries moved from empirical forecasts of the UV Index to

using RT models that can be very specific or can make

several assumptions to increase output speed. These

models provide the UV irradiances at the surface under

clear skies with known column ozone amount and aerosol

type and content. The difficulty in the forecast of UV

Indices lies in the forecast of the ozone amount and the

transmission of UV radiation through the atmosphere in

the presence of clouds and aerosols. Our inability to estimate accurately the aerosols at a specific location introduces further uncertainty in UV Index calculations, and a

recent study by Krzyscin et al. (2001) has addressed this

issue. Comparisons of model predictions with measured

UV Indices under cloudless conditions (De Backer et al.,

2001) showed that models generally overestimate the UV

Indices, owing to lack of aerosol information. Inclusion

of clouds in UV forecasts is a double problem, because

clouds must be accurately forecast and their transmittance

properties must be accurately known.

There have been two approaches to address the

cloud problem. In the absence of a comprehensive

weather forecasting model, empirical relations between

forecast cloud types and observed UV amounts provide

the necessary means of producing a UV Index forecast.

These relations may include variable amounts of aerosols

and the inherent elevation of the site, thus making the relationships unique to a particular site. The second approach

is to use a numerical weather prediction model (NWP)

such as that from National Centers for Environmental

Prediction (NCEP), the United Kingdom Meteorological

Office (UKMO), or the European Centre for MediumRange Weather Forecasts (ECMWF). These models provide the necessary cloud parameters, shortwave radiation

information, and snow cover (i.e., albedo) information to

derive a better estimate of the UV Index forecast.

Research must be conducted to make sure that the NWP’s

cloud fields and transmittances are validated by observations.

The forecasts of ozone used to produce the UV

Index forecasts have been determined using two methods:

persistence of satellite-observed ozone amounts, and

using NWP meteorological variables to make ozone forecasts via statistical regressions (Bais et al., 1998b; Plets

and Vynckier, 2000). The former procedure works well

for one-day forecasts year round in the tropics and during

summer poleward of the tropics. The latter method is

required when ozone variability is large and statistical

SURFACE ULTRAVIOLET RADIATION

5.10
SURFACE ULTRAVIOLET RADIATION

5.11

correlations between ozone variations and variations in

the meteorological parameters are greater.

Recently, forecasting centers (e.g., NWP mentioned above and the Koninklijk Nederlands Meteorologisch Instituut (KNMI)) have been assimilating ozone

into their forecast models. These models require a much

better depiction of ozone in the atmosphere for their radiation calculations. At NCEP, the Solar Backscatter

Ultraviolet Spectrometer (SBUV2) ozone profile and total

column information is assimilated into the global NWP

model. Accurate forecasts of the global ozone field are

available out to 5 days. Ozone forecasts are useful for

forecasting UV, as well as for the a priori ozone estimates

(total and profile) required by ozone retrieval algorithms.

The quality of this a priori value greatly affects the derived

value (Bhartia et al., 1996). Thus positive feedback

between observation and model forecast is created.

5.2.3 Remaining Questions and

Uncertainties

Remaining uncertainties in our understanding of

the radiative transfer processes that define surface UV

radiation include the following:

• A lack of knowledge of aerosol absorbing and scattering processes that results in observed differences

between some of the satellite estimates and groundbased measurements of surface UV irradiance.

• A lack of knowledge of aerosol optical depths and

aerosol single scattering albedo values on a global

scale. Differences in these variables from one geographic location to another may lead to differences in

the bias between ground-based measurements and

satellite estimates from one site to another.

• The extent to which inhomogeneities of terrain, surface reflectivity, or persistent geophysical features

surrounding an individual site influence groundbased measurements. The inhomogeneities also

affect how well a single point measurement is representative of an extended region (Chubarova et al.,

2002; Fioletov et al., 2002). These uncertainties, plus

the fact that the geographic distribution of groundbased sites is not uniform, make the determination of

global UV climatologies and long-term trends solely

from ground-based networks a difficult, if not impossible, task.

• The combined effects of clouds and surface albedo in

the derivation of satellite estimates of surface UV.

• The combined effects of clouds, aerosols, and tropospheric absorbing gases on ground-based and satellite measurements. Some of these processes have

been established in case studies, but their importance

is not quantified on global and regional scales.

5.3 AVAILABLE RESOURCES FOR STUDYING

SURFACE UV RADIATION

Surface UV radiation is studied by using several

types of resources that have been developed and applied

by many research groups. These resources include

ground-based instruments, RT models, satellite instruments, and statistical models. Ground-based instruments

that measure surface UV irradiance include spectroradiometers, broadband radiometers, and narrowband multifilter instruments. RT models calculate surface UV using

a model atmosphere with the solar spectrum and several

scattering and absorbing geophysical parameters as input.

Satellite instruments measure geophysical variables and

use RT models to calculate surface UV irradiance on a

global scale. Statistical models are used to determine

dependencies of surface UV irradiance on scattering and

absorbing variables.

5.3.1 Ground-Based Measurements

The measurement of solar UV radiation received at

the Earth’s surface is technically demanding. Usually, the

quantity measured is the irradiance on a horizontal surface using a detector with a cosine response. This means,

for example, that radiation at 60∞ from the zenith has only

half the weighting of that from the zenith direction.

Three categories of UV sensors are in widespread

usage: (1) spectroradiometers designed to measure the

spectrum of UV at UV-A (315-400 nm) and UV-B (280-

315 nm) wavelengths at spectral resolutions of 1 nm or

better, (2) broadband sensors designed to measure biologically weighted UV irradiance, in most cases the erythemally weighted (or “sunburning”) UV irradiance, and (3)

multifilter instruments where the irradiance is measured

through several narrowband (~2- to 10-nm bandwidth)

filters and the full spectrum can be recovered using RT

models.

5.3.1.1 SPECTRORADIOMETERS

Since the previous Assessment (WMO, 1999) there

have been a number of developments on UV spectroradiometers with the introduction of new instrument types

and improvements of existing UV instrumentation. A

dual-prism spectrograph (UV-rotating shadowband spectroradiometer) that makes continuous and nearly simultaneous spectral measurements of direct, diffuse, and total

horizontal irradiance using an array detector has recently
SURFACE ULTRAVIOLET RADIATION

5.12

been introduced (Harrison et al., 1999; Lantz et al., 2002).

Compared with conventional scanning spectroradiometers, these array instruments measure spectra much faster,

thus creating new opportunities to investigate fastchanging atmospheric variables such as clouds. Also a

high-resolution (0.1 nm) 1-m Czerny-Turner double

grating spectroradiometer with 10−10 out-of-band light

rejection has been developed and is in operation in the

U.S. Department of Agriculture (USDA) monitoring network (Bigelow et al., 1998; Lantz et al., 2002).

The need to correct for the cosine error (McKenzie

et al., 1992; Seckmeyer and Bernhard, 1993; Gröbner et

al., 1996) is now more widely recognized. More spectroradiometers have been equipped with diffusers that

represent the cosine response better than previously used

collectors (Bernhard and Seckmeyer, 1997; Bais et al.,

2001a). In addition, methods to correct measurements

made by instruments with known angular response have

been developed and applied to existing data records (Bais

et al., 1998a; Fioletov et al., 2002).

Wavelength shifts in measured spectra can contribute significantly to the overall error budget, particularly at wavelengths less than 300 nm where the spectral

gradient is large from absorption by ozone. Emphasis has

been given in recent years to develop improved wavelength drives with increased thermal stability and reproducibility (e.g., Gröbner et al., 1998). With the implementation of these developments, the wavelength scale of

many spectroradiometers can be aligned to within ±0.02

nm, which is acceptable for long-term trend detection.

Also methods for post-correction of wavelength shift by

correlating the Fraunhofer structure in measured spectra

with the same structure in reference spectra have been further refined and tested (Bais et al., 2001a; De la Casinière

et al., 2001).

The temperature dependence of the responsivity of

Brewer instruments in the U.S. Environmental Protection Agency (EPA) network has been characterized

(Weatherhead et al., 2001). It was shown that temperature dependence varies from instrument to instrument and

can result in errors up to ±10% in some cases. However,

with knowledge of the temperature dependence it is possible to partly correct existing data records.

A comprehensive list of spectroradiometer specifications based on requirements of UV research is given by

Seckmeyer et al. (2001). This report contains guidelines

for instrument characterization such asspectralsensitivity,

stray light determination, wavelength alignment, angular

response, and other parameters that influence the quality

of data measured by spectral UV instruments. The dominating factorsin the uncertainty budget ofspectroradiometers are uncertainties related to the radiometric calibration

(e.g., uncertainty of calibration standards), deviation from

the ideal angular response (the cosine error), instrument

drift, and wavelength misalignment (Bernhard and

Seckmeyer, 1999). Uncertainties arising from an instrument’s spectral bandpass (the slit function) and stray light,

which were historically a problem, are of less importance

with state-of-the-art instruments that use high-quality

double monochromators and a resolution smaller than 1

nm.

5.3.1.2 BROADBAND FILTER RADIOMETERS

Instruments that measure irradiance over a wide

wavelength range (>10 nm) are called broadband instruments. Many of these instruments are designed to measure

the erythemally weighted irradiance as defined by the

Commission Internationale de l’Éclairage (CIE)

(McKinlay and Diffey, 1987). From measurements of the

erythemally weighted irradiance, the UV Index (WMO,

1994) can be directly calculated. Since the previous

Assessment (WMO, 1999) research efforts have focused

on the calibration of these instruments and the analysis of

data, with little instrument development.

No broadband instrument precisely matches the

erythemal spectrum. Examinations of broadband instruments by independent laboratories have revealed that variation of individual instruments from the specifications

offered by manufacturers can result in calibration errors

between 10 and 20% (Bodhaine et al., 1998; Leszczynski

et al., 1998; Landelius and Josefsson, 2000). The conversion from individual detector-based units to standardized

units (e.g., the UV Index) depends on the wavelength

dependence of the radiation, and at the Earth’s surface this

dependence is characterized mainly by the SZA and total

ozone. Thus, the absolute calibration of a broadband

detector depends on these parameters. Often this dependence is determined using RT model results to convert from

detector weighted units to erythemally weighted units or

by comparing measurements of a broadband detector for

a considerably long time with a co-located spectroradiometer (Mayer and Seckmeyer, 1996; Blumthaler,

1997; Bodhaine et al., 1998; Lantz et al., 1999).

Significant changes in the responsivity of broadband detectors have been noted during routine operation

(Weatherhead et al., 1997; Silbernagl and Blumthaler,

1998; Borkowski, 2000). These changes are difficult to

detect without careful examination of the instruments

(Lantz et al., 1999; di Sarra et al., 2002b). Furthermore,

it was found that changes of internal relative humidity or

temperature are responsible for short-term variations of

sensitivity by more than 10% in some instruments (Huber

et al., 2002). Therefore, stability for these instruments (as
SURFACE ULTRAVIOLET RADIATION

5.13

well as for other UV instrument types) should be verified

by careful, periodic characterization and calibration.

5.3.1.3 NARROWBAND MULTIFILTER RADIOMETERS

Narrowband multifilter radiometers are less expensive and require less maintenance than spectroradiometers, yet, if well calibrated, can produce valuable information. Their bandwidths range from about 2 nm to 10 nm,

and some are equipped with an automated shadowband to

retrieve direct Sun irradiance. As with any radiometric

measurement, regular calibrations are essential. Recent

repeat measurements of the spectral response of 34 multifilter radiometers before and after 6 to 12 months of field

use revealed mean spectral shifts of less than 0.04 nm for

all channels (Gao et al., 2001). Multifilter radiometers

measure all wavelengths (typically from four to seven

channels) in less than 1 second as opposed to scanning

spectroradiometers that usually require several minutes to

complete a scan. This rapid and nearly simultaneous sampling of all wavelengths is useful for studying variations

in UV irradiance caused by rapidly changing cloud and

aerosol conditions. Radiometers that measure direct Sun

can retrieve aerosol optical depth (Wenny et al., 2001;

Slusser et al., 2002) as well as track the instrument’s radiometric stability in the field (Bigelow and Slusser, 2000).

Radiometers that measure global irradiance only, as well

as those measuring global and direct Sun irradiances, can

retrieve column ozone (Slusser et al., 1999; Gao et al.,

2001). Several methods have been successfully implemented to construct a complete spectrum from multifilter

measurements (Fuenzalida, 1998; Min and Harrison,

1998). Combining multifilter measurements with concurrent measurements of a scanning radiometer allows

the derivation of spectra at high temporal resolution, thus

blending the strengths of both instrument types (Thorseth

and Kjeldstad, 1999); however, rigorous intercomparisons

with spectroradiometers are still sparse.

5.3.1.4 DATA QUALITY AND DATABASES

Scientific objectives, such as the detection of trends

in UV radiation and investigations into radiative transfer

processes, require UV data of known quality that are

attainable only with carefully maintained and wellcharacterized instruments. The improvement of data

quality has remained a vital goal in recent years, and

progress has been made toward achieving this since the

previous Assessment.

Many new stations have started regular spectral UV

observations since the previous Assessment. However,

the geographic distribution of these stations is weighted

more toward continents in the Northern Hemisphere and

less over oceans and large regions of the Southern

Hemisphere. The existing nonuniform distribution is not

adequate to make measurements of long-term trends that

are representative of global coverage.

Currently there are two major databases where

interested scientists can obtain data: the World Ozone and

Ultraviolet Radiation Data Centre (WOUDC) operated

by the Meteorological Service of Canada (MSC) in

Toronto, Canada, and the European UltraViolet DataBase

(EUVDB) in Helsinki, Finland (called Scientific UV Data

Management (SUVDAMA) in the previous Assessment).

Appendix 5A gives a list of stations and data that are currently available. Since the contents of these databases are

rapidly changing, the interested reader is referred to the

database websites given in Appendix 5A as well as other

websites listed in Appendix 5B.

Most of the spectroradiometers deployed today are

radiometrically calibrated with tungsten halogen standard

lamps that are traceable to standards maintained by

national standards laboratories. The calibration uncertainty of these lamps is still one of the most prominent

sources of error in solar UV radiometry. Standard lamps

can abruptly change by up to 1% in the radiation output at

unpredictable intervals (Bernhard and Seckmeyer, 1999).

In addition, irreversible changes of 10% or more have

been seen by several investigators. These are connected

to visible changes in the lamp’s filaments. Such lamps can

no longer be used for calibration purposes. Lamp comparisons further suggest that even standards of the same

calibration laboratory may disagree with each other

beyond their stated accuracy, and deviations exceeding

4% have been observed at UV wavelengths (Bernhard and

Seckmeyer, 1999; Kiedron et al., 1999). National standards laboratories usually provide calibration points in

increments of 5 to 50 nm, and interpolation to wavelengths

between these points may lead to errors of the order of

0.5 - 1%. New interpolation procedures have recently

been developed to reduce this effect (Huang et al., 1998).

In order to diminish uncertainties related to transportation

of radiometers between calibration in the laboratory and

deployment on site, new field calibration units have been

developed in recent years (Seckmeyer et al., 1996, 1998;

Early et al., 1998; Bais et al., 2001b).

Assessment of the comparability of lamp standards

currently used at different stations in Europe has been initiated by the Joint Research Centre (JRC) of the European

Commission. A set of carefully selected and seasoned

lamps is circulated among the nine participating institutes,

which report their spectral irradiance output based on

comparative measurements with the local calibration standards. The results of the first round indicate differences

of up to 9% between the calibration standards of different
SURFACE ULTRAVIOLET RADIATION

5.14

laboratories (Gröbner et al., 2002). The average root mean

square (rms) difference of all the lamps to the mean is 2.0,

1.8, and 1.7% at 300, 330, and 360 nm, respectively, which

is within the expected uncertainty estimates of the lamp

calibration certificates. A series of measurements with

the same stationary spectroradiometer of JRC proved that

the lamps were stable to within 0.6% during the entire

period of the exercise, giving confidence to the differences encountered and to the feasibility of the method.

As an alternative to standard lamps, it has been proposed that the Sun can be used as a long-term reference to

monitor instrument stability and as a calibration source,

since its irradiance is stable to within ±0.5% at UV-A and

UV-B wavelengths. This is done with the Langley plot

technique using instruments that are able to perform both

direct and global measurements at clean sites where the

diurnal variability of ozone and aerosol extinction is small

(Bais, 1997; Slusser et al., 2000; Gröbner and Kerr, 2001;

Kerr, 2002). Ground-based measurements of direct solar

irradiance (using absolute lamp-based calibrations) are

extrapolated to zero air mass and compared with an extraterrestrial solar spectrum (e.g., Cebula et al., 1996). It has

been shown that lamp-based Langley extrapolation measurements of the solar spectrum performed with a Brewer

spectrophotometer under ideal conditions at Mauna Loa

Observatory (3397-m altitude) are consistent with satellite measurements to the 2-3% level for wavelengths

greater than 305 nm (Gröbner and Kerr, 2001). A similar

comparison of Langley and lamp-based calibrations was

performed at the same observatory with a UV multifilter

rotating shadowband radiometer. Differences between 0

and 7% were found, with largest deviations at the shortest

wavelengths (Slusser et al., 2000), and drifts in sensitivity

of the instruments of the order of 1% per year were

detected (Bigelow and Slusser, 2000). An important prerequisite for these extrapolation measurements is that

atmospheric conditions remain constant. Therefore such

calibrations can only be performed at very clean sites,

where aerosols do not influence UV irradiance. Also reliable calibrations cannot be made at wavelengths less than

305 nm, where the signal is strongly attenuated by ozone.

New techniques have been pursued to uncover

errors in spectral measurements. For example, analyzing

ratios of spectra sampled throughout a day enables the

detection of errors in the data acquisition of instruments

and the determination of wavelength shifts (Bernhard et

al., 1998; Seckmeyer, 2000). Standardized methods to

mark data with reduced accuracy or to flag scanning spectral UV measurements that are distorted by clouds have

been developed (Vasaras et al., 2001). The WOUDC now

flags data before adding them to the database, and other

data centers are currently developing data-flagging algorithms for the detection of spectral distortion, wavelength

error, and the presence of clouds.

Since the previous Assessment, more systematic

work to determine the uncertainty of UV measurements

has been completed. Bernhard and Seckmeyer (1999)

present a general procedure to calculate the uncertainty

budget of spectroradiometers depending on instrument

specifications, calibration uncertainties, and atmospheric

conditions. The method can be applied to most types of

spectroradiometers deployed worldwide. For wellcharacterized spectroradiometers, erythemal irradiance

can be measured within an uncertainty of ±6% (±2s). A

substantial reduction in this uncertainty would require

more accurate calibration sources and improved methods

to correct for instrument drifts. Below 300 nm, the most

important sources of error are typically wavelength misalignment and photon noise. Thorough uncertainty estimates for actinic flux spectroradiometers (i.e., instruments

with isotropic rather than cosine-weighted angular

response) have also been performed recently

(Hofzumahaus et al., 1999). According to Hofzumahaus

et al. (1999), the total uncertainty of photolysis frequencies due to uncertainties in the measurement of the spectral actinic UV flux lies in the range of 5-7% and is dominated by the ±4% (±2s) uncertainty of the irradiance standard. This result is similar to the uncertainty estimate for

global irradiance measurements.

In 1994, WMO established a scientific advisory

group to provide guidance for UV measurements performed within the Global Atmosphere Watch program.

Guidelines for site quality control of UV monitoring

(Webb et al., 1998) and recommended specifications for

spectroradiometers measuring solar ultraviolet radiation

(Seckmeyer et al., 2001) have been drafted. Both publications are considered as working documents that will

evolve when new technologies or new objectives for UV

spectroradiometry emerge.

5.3.1.5 INSTRUMENT INTERCOMPARISONS

Since the beginning of the 1990s, periodic intercomparisons of spectroradiometers from different organizations have become an international practice for

assessing the quality of UV radiation measurements.

Although the stability and the absolute accuracy of individual instruments can be sufficiently maintained using

various calibration methods, intercomparison with other

independently maintained instruments is an important

method that provides uniform quality assurance and data

quality. Intercomparisons also provide the opportunity

for effective interaction between participating scientists

in order to exchange new ideas and findings.
Results of three intercomparisons have been

reported since the previous Assessment. The SUSPEN

(Standardization of Ultraviolet Spectroradiometry in

Preparation of a European Network) intercomparison held

in Greece in July 1997 (Bais et al., 2001a) included 19

spectroradiometers from 15 countries. In August 1997 an

intercomparison of 10 spectroradiometers operated in

Germany (Seckmeyer et al., 1998) took place. In

September 1997 the fourth North American Intercomparison of Ultraviolet Monitoring Spectroradiometers

(including narrowband filter radiometers) was held near

Boulder, Colorado. It included 11 instruments from 9 U.S.

agencies (Lantz et al., 2002). Results of these intercomparisons showed that agreement of about ±5-6% (1s) was

typically achieved, with some dependence on wavelength,

zenith angle, and observing conditions. Agreement

improves with the application of corrections (e.g., wavelength shift and cosine error) to the data and is generally

better at longer wavelengths.

The descriptions above represent gross summaries

only. The behavior of the individual instruments can only

be judged from more detailed information found in the

references. Although the intercomparisons have shown

that a considerable number of instruments still show large

deviations, overall improvement has been achieved as a

result of the knowledge gained through the intercomparison exercises.

Previous comparisons and assessments of the operation of broadband instruments have shown the need to

examine their performance regularly (e.g., Johnsen and

Moan, 1991; DeLuisi et al., 1992; Weatherhead et al.,

1997). In this context an intercomparison campaign

involving 39 erythemal radiometers and two wellcalibrated and maintained spectroradiometers was organized in 1999 at the University of Thessaloniki, Greece

(Bais et al., 2001b). The campaign was combined with

laboratory investigation of the spectral and angular

response of the participating instruments. From the comparisons of broadband with the spectroradiometric measurements, new calibration factors were derived, with

differences from the original values ranging between

–10% and +25%. For about half of the instruments the

differences were smaller than about ± 10%, but only a few

agreed with the spectroradiometer to better than ± 5%.

These deviations depend on the time elapsed between the

two subsequent calibration checks and on the individual

characteristics of each instrument. The large deviations

found for many of the instruments suggest that the calibration of broadband detectors should generally be

checked more frequently, either in comparison with a

spectroradiometer, or against another instrument of

similar type.

5.3.2 Radiative Transfer Models and Model

Validation

Radiative transfer (RT) models allow the calculation of radiation fluxes (direct, global, sky radiance distribution) as functions of geophysical variables and SZA.

Most common models are one-dimensional (planeparallel or pseudo-spherical) with scattering on spherical

particles and polarization effects neglected. Several of

these models are freely available for public use (Internet

sites given in Section 5B.3 of Appendix 5B): the “library

for Radiative transfer” (libRadtran, formerly known as

uvspec) (Mayer et al., 1997; Kylling et al., 1998); the

Tropospheric Ultraviolet Visible (TUV) model

(Madronich and Flocke, 1997); the Santa Barbara

Discrete-ordinate Atmospheric Radiative Transfer

(SBDART) model (Ricchiazzi et al., 1998); the Streamer

model (Key, 1999); and the System for Transfer of

Atmospheric Radiation (STAR) model (Ruggaber et al.,

1993, 1994). However, these one-dimensional models

cannot be used to study situations with nonuniform spatial distributions of scattering or absorbing parameters.

Comparison of measurements of spectral UV surface fluxes with numerical RT simulations greatly facilitates the interpretation of measurements. Also, the quality

and internal consistency of measurement methods may be

checked using numerical simulations. In addition, RT calculations are required for estimating surface UV radiation

from satellite data and for making forecasts of spectral

surface UV radiation and the UV Index. In order to use

RT simulations for these purposes, the validity of the simulations must be assured. However, validation of UV RT

models for real atmospheric cases is a complicated task.

At best the validation includes a model intercomparison

in combination with observations for some well-defined

cases. This combined approach has been pursued in the

Scientific UV Data Management (SUVDAMA) project

(van Weele et al., 2000).

The SUVDAMA model intercomparison (with 12

participating codes) established benchmarks for six real

cloud-free atmospheric cases with different ozone column,

air density profile, SZA, aerosol loading, surface albedo,

and location. The benchmarks give the spectral surface

UV irradiance between 295 and 400 nm with 0.5-nm steps

and with a standard deviation of 2% between the modeled

spectral irradiances. For cases with the Sun close to the

horizon and for the shortest wavelengths, the relative standard deviations are somewhat higher (5%), but then irradiance is very small. Remaining differences between

model results relate to the translation of the ancillary data

and other fixed input parameters into the optical properties that are needed in the RT algorithms. Examples

SURFACE ULTRAVIOLET RADIATION

5.15
SURFACE ULTRAVIOLET RADIATION

5.16

include the parameterization of aerosol optical properties,

interpolation of the ozone absorption cross section as a

function of wavelength and temperature, and the discretization of the atmosphere into vertical homogeneous

layers. Comparison of the benchmarks with the observations showed differences, within ±13% over the whole

spectral UV range for four out of six cases, mainly due to

uncertainty in the input parameters and assumptions on

instrument characteristics such as slit function and wavelength alignment.

Atmospheric and ground parameters are usually

horizontally inhomogeneous, especially for cloudy cases.

These nonuniformities must be addressed by using threedimensional (3-D) models, which require significant computational time and are therefore restricted to a subset of

conditions only. In recent years, 3-D RT models have

become available, which use either the Monte Carlo

method (e.g., Cahalan et al., 1994) or the spherical harmonics discrete ordinate method (Mueller and Crosbie,

1997; Evans, 1998). These models allow the study of

effects of three-dimensional clouds (O’Hirok and Gautier,

1998a, b; Marshak et al., 1999; Degünther and Meerkötter,

2000a; Meerkötter and Degünther, 2001) or of inhomogeneous surface albedo (Degünther et al., 1998; Degünther

and Meerkötter, 2000b; Kylling et al., 2000).

The first intercomparison of 3-D radiation codes

has been performed. Differences between the 3-D model

calculations are in the range ±5-10%, which is comparable with the uncertainties typically associated with UV

measurements. However, such a complete description of

the atmosphere is usually not available when UV measurements are to be compared with results of model calculations. Therefore, the uncertainty of the calculated irradiance is much higher, even under clear-sky conditions

(Weihs and Webb, 1997a, b; Schwander et al., 1997). For

well-characterized conditions, agreement to within about

±5-10% can be obtained between observations and model

calculations. When some input parameters are not known,

differences higher than ±20% are found (Koepke et al.,

1998; De Backer et al., 2001).

5.3.3 Satellite Estimates and Validation

The understanding of RT processes forms the scientific basis for estimating surface UV irradiance from

satellite measurements. Since the previous Assessment

(WMO, 1999), advances have been made in our understanding and application of the required processes. The

Total Ozone Mapping Spectrometer (TOMS) UV algorithm (Eck et al., 1995; Herman et al., 1996; Krotkov et

al., 1998) has been modified to improve cloud attenuation

estimates and to introduce the enhancement of surface

irradiance caused by the high albedo of snow (Herman et

al., 1999; Krotkov et al., 2001, 2002). The method has

been adapted to the Global Ozone Monitoring Experiment

(GOME), the European UV spectrometer on the second

European Remote Sensing Satellite-2 (ERS-2) (Peeters et

al., 1998). Li et al. (2000) have proposed an alternative

parameterization for deriving surface irradiance from

TOMS total ozone and TOMS 380-nm reflectivity or

Advanced Very High Resolution Radiometer (AVHRR)

visible data. Mayer et al. (1998a) have modified the

TOMS UV algorithm to estimate the surface actinic flux

giving photolysis rates for atmospheric chemistry models.

Polar-orbiting UV spectrometers (e.g., TOMS,

GOME) are able to make measurements of atmospheric

properties at a given location once per day. This together

with a relatively coarse spatial resolution affects their

ability to estimate the total daily exposure to UV radiation because of the high temporal and spatial variability

of cloud cover. It has been shown that for accurate

retrievals, several cloud images per day are required

(Martin et al., 2000; Meerkötter and Bugliaro, 2002;

Verdebout and Vogt, 2002). Different methods have been

proposed that draw cloud transmission information from

other satellites or datasets with higher spatial and/or temporal resolution. Lubin et al. (1998) used Earth Radiation

Budget Experiment (ERBE) hourly cloud and surface

albedo data resolved at 100 km to determine a 5-year climatology; Matthijsen et al. (2000) generated a multiyear

dataset over Europe using the International Satellite Cloud

Climatology Project (ISCCP) dataset (~15 km, every 3

hours); and Verdebout (2000) retrieved cloud optical

thickness and surface albedo from METEOSAT (~5 km

over Europe, half-hourly) to generate UV maps over

Europe. AVHRR data (between one and five images per

day with ~1-km resolution) have been used to map the

surface UV radiation over Antarctica (Lubin et al., 1994),

the Moscow region, Russia (Rublev et al., 1997), and

parts of Europe (Meerkötter et al., 1997) as shown in

Figure 5-1 (from Simon et al., 2000). These methods use

the TOMS, GOME, or Television Infrared Observation

Satellite (TIROS) Operational Vertical Sounder (TOVS)

derived total column ozone and various ancillary information such as digital elevation models, aerosol climatologies, or gridded surface visibility observations as

inputs to RT models that estimate surface UV irradiance.

Accurate knowledge of cloud transmittance is critical for satellite UV estimates. The TOMS and GOME

instruments can measure UV cloud reflectivity, R, and estimate cloud transmission, CT, by the simple expression CT ~

1 – R, with a correction for ground reflectivity (Eck et al.,

1995; Herman and Celarier, 1997; Herman et al., 2001a;

Krotkov et al., 2001). Alternately, CT for plane-parallel
SURFACE ULTRAVIOLET RADIATION

5.17 ) N g e d( e d ut i t a Ltude (deg N) L

Longitude (deg E) m above MSL m a b o v e M S L ) N g e d( e d ut i t a Ltude (deg N) L

Longitude (deg E) Cloud Optical Depth C l o u d O p t i c a l D e p t h ) N g e d( e d ut i t a Ltude (deg N) L

Longitude (deg E) o D U ) N g e d( e d ut i t a Ltude (deg N) L

Longitude (deg E) m W / m 2 (a)

(c)

(b)

(d)

Figure 5-1. A region in Central Europe including the Alps: (a) surface elevation in meters above mean sea level, (b) horizontal distribution of the

GOME (Global Ozone Monitoring Experiment) total ozone column amount in Dobson units (DU) on 15 April 1997, (c) horizontal distribution of cloudoptical depth as derived from NOAA/AVHRR (Advanced Very High Resolution Radiometer) data on 15 April 1997, and (d) high-resolution map ofthe erythemal dose rates in milliwatts per square meter (mW m-2). From Simon et al. (2000).
SURFACE ULTRAVIOLET RADIATION

5.18

cloud can be estimated by radiative transfer inversion of

the satellite-measured “top of atmosphere radiance.” The

latter approach is applicable to UV, visible, and infrared

sensing instruments, since the CT spectral dependence is

included in the RT model (Meerkötter et al., 1997; Krotkov

et al., 2001; Verdebout, 2000). There have been a number

of studies where satellite estimates of CT compared with

ground-based measurements often show good agreement,

but reveal systematic errors in the presence of snow

(Kalliskota et al., 2000; Arola et al., 2002). A method to

improve satellite cloud reflectivity estimates with snow on

the ground has been developed by including snow depth

information (Krotkov et al., 2002), and this method is to be

considered for future satellite algorithms.

Under cloud-free conditions the accuracy of satellite UV data is limited mainly by an imperfect knowledge

of the highly variable aerosol properties. TOMS UV twochannel reflectivity data are used to correct for absorbing

aerosols at a known altitude by forming the aerosol index

(AI; Krotkov et al., 1998). The effect of aerosols on satellite UV estimates over large areas is largest in tropical

regions where there are major dust plumes (e.g., the

Saharan plume) and smoke plumes (e.g., Africa and South

America) from biomass burning. There the reductions in

surface UV irradiance are frequently as much as 50%

(Herman et al., 1999). However, the TOMS AI technique

does not have the sensitivity to detect absorbing or nonabsorbing aerosols close to the ground that are often

present in urban atmospheres (Dickerson et al., 1997;

Torres et al., 1998; Herman et al., 1999; Jacobson, 1999).

When the aerosol effects are estimated using measured

optical depths and single scattering albedos between 0.95

and 0.98, the clear-sky differences are usually reduced to

within the instrument uncertainty. However, when nearsurface atmospheric visibility data are used, UV irradiance reductions as large as 10% are found in some parts

of Europe (corresponding to visibility values of ~10 km)

(Verdebout, 2000), and even higher reduction could be

expected in some cases. Simultaneous measurements of

aerosol optical properties and surface UV irradiance are

required to quantify local satellite biases for such areas.

The aerosol problem is smaller at midlatitudes in the

Southern Hemisphere because of clearer air compared

with the same latitudes in the Northern Hemisphere

(McKenzie et al., 2001a).

The accuracy and precision of satellite estimates

have been assessed through comparisons between satellite and ground-based data and are reported either in

papers describing the satellite-based methods or in dedicated publications (Kalliskota et al., 2000; Wang et al.,

2000; Herman et al., 2001b; McKenzie et al., 2001a; Arola

et al., 2002; Chubarova et al., 2002; Fioletov et al., 2002;

Slusser et al., 2002, Wuttke et al., 2002). The observed

differences of noontime irradiance values between

ground-based measurements and TOMS satellite estimates can be ±20-30% rms and are caused by temporal

and spatial cloud variation. However, the rms differences

are reduced for daily integrals averaged over longer

periods (e.g., ±4% for May, June, July, and August

monthly mean values at Toronto; Fioletov et al., 2002).

McKenzie et al. (2001a) compared ground-based measurements from cross-calibrated spectrometers with UV

estimated from TOMS instruments over several years at

four midlatitude sites. There is reasonable agreement in

the day-to-day variability of UV derived by both methods;

however, the TOMS data overestimate monthly erythemal

doses by about 15% in Toronto, 24% in Thessaloniki, and

38% in Garmisch-Partenkirchen. At the pristine Southern

Hemisphere site, Lauder, New Zealand, the average agreement is within 3%. Fioletov et al. (2002) reported similar

comparisons at 10 sites in Canada. At 9 of the 10 sites

TOMS overestimates the ground-based measurements

from 7 to 16%. The only site with agreement to within

2% is at Saturna, located on the Canadian west coast where

there is predominantly clean maritime air. Chubarova et

al. (2002) showed that TOMS estimates are about 10%

larger than a long-term (1979-2000) record of broadband

measurements made at Moscow State University. The

reasons for the large differences at northern midlatitudes

may be attributed to absorption in the lower troposphere

by aerosols and/or other pollutants that are not properly

considered in the TOMS retrievals.

Part of the observed differences reflects the fact

that the estimates and measurements are not of the same

thing. There is always a mismatch between satellite-based

and ground-based data acquisition. The satellite-derived

value represents a single large-area average compared

with the near-continuous local irradiance measured by a

ground-based instrument. Lubin et al. (1998) have shown

that daily doses can be estimated from satellite data

obtained only at local noon with knowledge of the course

of the SZA throughout the day. For all cloud cases Martin

et al. (2000) have shown that the uncertainty in daily doses

reconstructed from a single ground-based near-noon

measurement is at least ±25%, and reduces to ~±5% for

monthly doses.

Another important consideration is that there are

site-specific biases caused by the persistence of local geophysical variables (McKenzie et al., 2001a, b; Chubarova

et al., 2002; Fioletov et al., 2002). These include variations in altitude around the site, the persistence of local

cloud patterns, or persistent patterns of nearby albedo.

The 38% bias between TOMS and GarmischPartenkirchen ground-based UV data (reported by
SURFACE ULTRAVIOLET RADIATION

5.19

McKenzie et al., 2001a) was not seen when METEOSAT

or AVHRR data were used (Arola et al., 2002). The algorithm developed by the Institut d’Aéronomie Spatiale de

Belgique (IASB) uses TOMS data, but shows a mean difference of 19% (Figure 5-2). The reason for the difference between the original TOMS algorithm and the IASB

algorithm is not known. The algorithm of JRC based on

METEOSAT and GOME data, and the algorithm of RIVM

(National Institute of Public Health and the Environment,

Netherlands) based on TOMS ozone and ISCCP data,

show a smaller difference between satellite estimates and

ground-based measurements at Garmisch-Partenkirchen

as well as other sites (Figure 5-2). This leads to the conclusion that some satellite algorithms give better estimates

of surface erythemal UV radiation than others. In addition there are systematic seasonal variations between

satellite and ground-based results that are not understood

yet. A novel method to understand such differences better

has been proposed by Wuttke et al. (2002). It has been

shown that the use of the full spectral information of both

the satellite algorithm and the ground-based measurements helps to identify the reasons for deviations. Such

an improved understanding has been used to develop an

improved algorithm that shows less deviation when compared with ground-based measurements.

5.3.4 Statistical Modeling

The main geophysical variables that affect surface

UV radiation are ozone, clouds, aerosols (both absorbing

and nonabsorbing), and surface albedo. Total ozone measurements and ancillary data have been used with measurements of surface UV irradiance to develop statistical

relationships that define the dependence of UV on the

scattering and absorbing variables. Previous applications

of statistical models include the UV Index forecast (e.g.,

Burrows et al., 1994) and the derivation of total ozone

from UV measurements (Fioletov et al., 1997).

Recent studies (Krzyscin, 1996; Fioletov et al.,

1997, 2001; McArthur et al., 1999; Bodeker et al., 2000;

Kaurola et al., 2000; Gantner et al., 2000; den Outer et al.,

2000; Diaz et al., 2000, 2002; Chubarova et al., 2002)

have shown that surface UV radiation can be estimated

by using measurements of total ozone (ground-based or

satellite) with other ground-based data (such as global

solar radiation measured with pyranometers). Total ozone

values determine the spectral absorption features in the

UV-B, and the ancillary data quantify the scattering

processes that have weak wavelength dependence (clouds,

aerosols, and surface albedo).

There have been a number of approaches for the

use of statistical models. In some of these models, daily

integrals of wavelength interval or erythemally weighted

UV irradiance are statistically related to measurements of

global radiation and other atmospheric variables (Diaz et

al., 2000, 2002; Gantner et al., 2000); in others (Kaurola

et al., 2000; den Outer et al., 2000) the daily integrals are

compared with model results. Relationships as functions

of total ozone, SZA, and other variables allow hour-byhour comparisons (Bodeker et al., 2000; Fioletov et al.,

2001). The neural network technique has also been

applied to datasets of total ozone, global pyranometer

information and other weather observations (Janouch,

2000; Schwander et al., 2002).

Good-quality spectral UV irradiance measurements

have been available for about 10 years. These data are

used with total ozone and ancillary data taken over recent

-40

-20

0

20

40

60

1 2 3 4 5 6 7 8 9 10 11 12

SATELLITE - GROUND-BASED UV (%)

MONTH

IASB JRC RIVM

mean diff(%) 19.094

rms diff(%) 22.492

-2.072

11.389

-4.051

17.136

IASB

JRC

RIVM

Figure 5-2. Percentage difference

between ground-based and satellite-based

monthly erythemal irradiance in 1997 predicted by three different satellite algorithms:

from the Institut d’Aéronomie Spatiale de

Belgique (IASB), the European Commission Joint Research Centre (JRC), and

the National Institute of Public Health

and the Environment of the Netherlands

(RIVM). The algorithm of IASB shows a

systematic bias with respect to groundbased data, whereas the algorithms of JRC

and RIVM show improved agreement for

the mean erythemal yearly dose. However,

the latter two algorithms show a seasonal

dependence, which is not well understood

yet. Adapted from Arola et al. (2002).
years to establish statistical relationships that determine

the dependence of ground-based spectral UV radiation on

total ozone and the ancillary data. The statistical relationships are then used with measurements of total ozone and

ancillary data to estimate UV radiation at other sites and

at times when spectral UV radiation was not measured.

These estimates can be extrapolated backward in time to

as early as the 1960s, prior to the onset of ozone depletion. Figure 5-3 shows an example of statistical model

results compared with ground-based spectral UV irradiance measurements and TOMS UV satellite estimates for

summer months at Toronto (Fioletov et al., 2002). This

figure demonstrates the good agreement between the

ground-based estimates and spectral measurements for

the overlap period (1989-1997). Also, the year-to-year

variability of the ground-based estimates extended backward is similar to that of the satellite estimates, although

there is a bias between satellite and ground-based results

as discussed earlier.

5.4 UV CLIMATOLOGY, TEMPORAL CHANGES,

AND TRENDS

The main objectives for taking long-term systematic measurements of surface UV radiation are to establish a global climatology of UV, both average and extreme

values, and to quantify any long-term changes that may

have occurred as a result of changes in stratospheric ozone

or other variables. This information is of interest to members of the “effects” community, who focus research

studies on the response of biological systems (including

human beings), materials, and pollution photochemistry

to changes in the UV environment. Some ground-based

spectral datasets are available to establish average values

and variability of UV at specific sites; however, these

datasets are relatively short in duration and have sparse

and nonuniform spatial coverage over the globe. They

are also representative of areas that are small in size and

defined by specific local conditions. Understanding and

quantifying the radiative transfer processes (Section 5.2)

has allowed the extension of these datasets, both in time

and space, with a good degree of certainty. Global coverage has been achieved by the use of satellite data, and

the temporal extension backward in time has been

achieved by using satellite measurements with RT models,

as well as ancillary ground-based data with RT and statistical models.

5.4.1 Ground-Based UV Measurements

Difficulties involved in the routine operation and

maintenance of accurately calibrated ground-based UV

instruments have limited the length of reliable data records

to about the past 10 years. It is recognized that 10-year

records are not of adequate length to carry out long-term

trend analyses (Weatherhead et al., 1998, 2000) because

of a combination of variability and autocorrelation in

datasets. However, some recent studies have used groundbased measurements to identify shorter term changes in

surface UV radiation (McKenzie et al., 1999; Frederick et

al., 2000; Zerefos, 2002). Also there has been some work

done in using the data to develop specific climatologies

of UV radiation at some sites (Ilyas et al., 1999; Cede et

al., 2002).

An increase in peak UV values in response to

decreasing ozone at Lauder, New Zealand, is shown in

Figure 5-4 (from McKenzie et al., 1999, 2000). The trend

has not continued in the last two austral summers, and

ozone amounts at this site have been slightly higher than

in the summer of 1998/99. Furthermore, both summers

were rather cloudy over the period that is most critical for

this analysis. It should be noted that year-to-year variability in cloud cover can have a significant effect even

for peak irradiances and that the analysis of peak values

cannot be automatically transferred to average values.

SURFACE ULTRAVIOLET RADIATION

5.20

1965 1970 1975 1980 1985 1990 1995 2000

3000

3200

3400

3600

3800

4000 Summer Mean CIE Irradiation (J m- 2)

Brewer measurements

Derived frompyranometer data

Derived fromTOMS

Figure 5-3. Time series of Toronto summertime

(May-August) mean daily ground-based and satellitebased erythemal irradiation estimates compared

with ground-based measurements. The estimated

values using ground-based total ozone and pyranometer data agree very well with the measurements

during the overlap period (1989-1997). The satellite

estimates follow the year-to-year variability of the

ground-based measurements; however, there is a

bias as discussed in the text. Similar year-to-year

variability is captured by both the satellite-based and

ground-based estimates extrapolated backward in

time. Adapted from Fioletov et al. (2002).
SURFACE ULTRAVIOLET RADIATION

5.21

Data analysis for the summer of 2000/01 has not yet been

finalized because of uncertainties in the National Institute

of Standards and Technology (NIST) irradiance scale provided by the irradiance standards laboratories involved.

It is likely that the entire dataset will eventually be reanalyzed taking historical changes in lamp irradiance scales

into account better. However, it is not expected that the

conclusions of the paper will change as a result of that

reanalysis.

Routine spectral ultraviolet irradiance measurements made at Reading, England, since 1993 (Bartlett and

Webb, 2000), along with total ozone measurements from

a nearby site, demonstrated the inverse relationship

between ozone and UV-B. Global radiation measurements

suggest that from 1993 to 1997 cloud conditions did not

change systematically. A decline in ozone of 5.9% and a

corresponding increase in erythemal UV of 4.3% were

seen, although these changes should not be considered

statistically significant trends. Also, solar UV irradiance

spectra (290-325 nm) have been measured at Rome and

Ispra, Italy, since 1992 (Casale et al., 2000). Correlation

coefficients between irradiance at 305 nm and total ozone

at Rome were –0.61 (SZA = 47°) and –0.75 (66°) and at

Ispra, –0.55 and –0.76. Seasonal analysis showed higher

negative correlation with short-term ozone changes

(related with weather patterns) during spring and winter,

while in summer the long-term changes (time period

greater than 2 years) and seasonal changes (between 2

years and 1 month) in ozone were dominant.

Time series of UV irradiance at 305 and 325 nm

for Thessaloniki, Greece, for 1990-1997 were reported in

the previous Assessment (Zerefos et al., 1997, 1998;

Herman and McKenzie et al., 1999) for low-cloudiness

( 2/8 cloud cover) conditions. These series have been

updated to 2001 (Zerefos, 2002) and are shown in Figure

5-5. Both ozone and sulfur dioxide (SO2) have decreased

during the 1990s (Zerefos et al., 1998). However, the

observed increase in irradiance of 16.3% per decade at

260

270

280

290

300

310

320

Ozone (DU)

78/79

79/8080/81

81/8282/8383/84

84/8585/86

86/8787/88

88/8989/90

90/9191/9292/93

93/9494/95

95/9696/9797/98

98/9999/00

Summer Year (December - February)

9.5

10.0

10.5

11.0

11.5

12.0

UV Index at noon

(b)

(a)

12

1

2

12

1

2

12

1

2

121

2

12

1

2

1

2

12

1

2

12

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121

2

121

2 12

1

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12

12

121

2

12

1

2

12

1

2

12

1

2

1990 1992 1994 1996 1998 2000 2002

YEAR

50

90

130

SOLAR IRRADIANCE (mW.m-2.nm-1) 1.5

3.0

4.5

300

350

400

TOTAL OZONE (m-atm-cm)

300

350

400

Irradiance at 305 nm

Irradiance at 325 nm

(O3: -1.3 %/decade; 305 nm: 20%/decade; 325 nm: 16.3%/decade)

Figure 5-4. (a) Mean total ozone and (b) noontime

UV Index at Lauder, New Zealand, for summers

(December to February) from 1978/79 to 1999/2000.

The solid line in (a) shows observed summertime

ozone values that have occurred since the 1970s,

and the solid line in (b) shows the expected values

of clear-sky UV. The symbols (from 1989/90 on)

show measured summertime values of ozone and

peak UV Index, both derived from spectral UV irradiance measurements. The plot also shows the values

from each contributing month, labeled by the month

number. Two-sigma error bars are shown for reference. Adapted from McKenzie et al. (1999, 2000).

Figure 5-5. Time series of low-cloudiness ( 2/8

cloud cover) monthly mean values of total ozone and

solar UV irradiance (63° SZA for 305 and 325 nm)

measured at Thessaloniki, Greece (40°N), between

1990 and 2001. These time series are extensions

of those reported in the previous Assessment

(Herman and McKenzie et al., 1999; Zerefos et al.,

1997). The observed temporal increase at 325 nm

is attributable to causes other than ozone, because

ozone absorbs weakly at this wavelength.
325 nm cannot be caused by the decrease of these gases,

since they both absorb weakly at this wavelength. The

increase must be caused by long-term changes in other

factors, such as aerosols that are associated with the

decrease in SO2 at the observing site.

The spectral UV irradiance records of the U.S.

National Science Foundation (NSF) UV radiation monitoring network operating at Ushuaia (Argentina) and

Palmer, McMurdo, and South Pole (Antarctica) have

increased in length and are now more than 10 years long.

Recent analyses of the records have been carried out

(Sobolev, 2000; Booth et al., 2001; Diaz et al., 2001).

Although the NSF dataset spans roughly 10 years, it is

still too short for statistically robust trend detection. At

all Antarctic sites, changes in erythemal and DNAweighted UV are generally positive for the months from

September to December, with largest increases typically

observed in November. However, trends in monthly averages of daily erythemal and DNA-weighted doses are not

significant at the 2-sigma level because of the large yearto-year variability in total column ozone, related to the

interannual variability of the ozone hole. Compared with

ozone variability, the year-to-year fluctuation in other

factors such as cloud cover is of minor importance for

interannual UV changes at Antarctic network sites.

The highest UV levels on record observed at

McMurdo, Palmer Station, and the South Pole by the NSF

network instruments occurred in the austral spring of

1998. UV levels in 1999, 2000, and 2001 were generally

lower. The record-size ozone hole in 2000 led to enhanced

UV levels at all austral network sites during September,

and some extreme events at Ushuaia during October. The

observed increases during September are small in absolute

terms because of the low solar elevations prevailing

during this month. At all austral network sites, UV values

in November 2000 were close to the minimum values on

record because of the early closing of the ozone hole.

Yearly maximum erythemal UV levels at austral

network locations were observed during recent years in

November and early December, 1 to 2 months after the

period of maximum ozone depletion. This pattern is

explained by the combination of smaller SZA later in the

year and the relatively low stratospheric ozone concentrations that still prevail in November. Figure 5-6 demonstrates that peak levels of daily erythemal UV dose in

November and December at austral sites exceed those seen

during the summer at San Diego (Booth et al., 2001). The

large enhancements in November and December suggest

that changes in the duration of the ozone hole have a

greater influence on biologically relevant UV levels than

does the depth of minimum total ozone typically observed

at the end of September and beginning of October (Diaz

et al., 1994).

5.4.2 Estimates from Satellite Data

The UV trends calculated from the Nimbus-7

TOMS satellite data (1979-1992) were reported in the previous Assessment (Herman et al., 1996; Herman and

McKenzie et al., 1999; Ziemke et al., 2000). Follow-on

TOMS missions (Meteor-3, Advanced Earth Observing

Satellite (ADEOS), and Earth Probe) have extended the

TOMS UV record, with an 18-month gap in 1995-1996

(Herman et al., 1999). Extensions of the trends calculated

on a global scale and presented in the previous Assessment

are not yet available.

SURFACE ULTRAVIOLET RADIATION

5.22

0

2

3

4

5

6

7

8

9

10

01-Jan 20-Feb 10-Apr 30-May 19-Jul 07-Sep 27-Oct 16-Dec

Date Erythemal Daily Dose (kJ/m2)

McMurdo

Palmer

South Pole

Ushuaia

San Diego

Barrow

Figure 5-6. The maximum daily

dose of erythemal irradiance (in

kilojoules per square meter)

measured at stations in the

National Science Foundation

(NSF) UV monitoring network.

Peak values occur at austral

sites in middle to late spring

when the combination of higher

Sun angles and low values of

total ozone still prevail from

an extended breakup of the

Antarctic ozone hole. The peak

values at these high-latitude

sites exceed those at significantly lower latitudes during

summer (e.g., San Diego).
SURFACE ULTRAVIOLET RADIATION

5.23

Both the TOMS-derived and TOMS/ERBE-derived

global UV climatologies have documented the seasonal

surface UV variability in different regions in the world

caused by cloudiness, ozone, UV-absorbing aerosols, and

surface albedo (Lubin and Jensen, 1995; Lubin et al.,

1998; Herman et al., 1999, 2001a, b). In addition, interannual variability associated with the quasi-biennial oscillation has been observed (Herman et al., 2000). These

datasets permit us to identify the areas and periods where

particularly high levels of UV radiation can be expected

in cases of low ozone events and low amounts of cloud

cover (e.g., parts of South America, Africa, and Australia).

Year-to-year variations of UV values in March are up to

±35% over Europe because of highly variable ozone and

cloudiness. There have also been changes in UV irradiance associated with long-term systematic changes in

cloudiness (Herman et al., 2001a), with large regions of

cloud-induced increases over Europe and parts of North

America, as well as some small but well-defined regions

of UV decrease over the oceans near South America and

Antarctica. This represents significant progress in understanding since the previous Assessment (WMO, 1999).

At that time the decrease of UV due to the increase in

cloudiness over Europe was identified from ground-based

measurements of total (pyranometric) radiation but was

not found in the satellite estimates.

Trends in snow cover and/or albedo can also affect

UV trends at high latitudes. The effects of these forcings

on UV trends are significant, but their relative contribution varies geographically and seasonally. Vinnikov et al.

(2002) found a statistically significant decrease in sea ice

cover over the Northern Hemisphere, whereas sea ice

cover has increased in the Southern Hemisphere. Thus it

is expected that UV irradiance has decreased in some areas

of the Northern Hemisphere and has increased in some

areas of the Southern Hemisphere. These conclusions can

be drawn from our understanding of the radiative transfer,

but ground-based observations are lacking, especially

over oceans. Satellite estimations of UV in these areas

and seasons are complicated by the difficulties in distinguishing between snow/ice and cloud albedo.

Whereas the effects of UV-absorbing aerosols on

surface UV irradiance have been investigated (Krotkov et

al., 1998; Herman et al., 1999), their effects on UVirradiance trends have not been estimated. In some situations, there are indications that aerosol effects are an

important source of error (McKenzie et al., 2001a). In the

tropics, there does not appear to be a long-term trend in

dust aerosol amounts, even though biomass burning in

South America has caused some increases in smoke

(Herman et al., 1997), which considerably absorbs UV

radiation. Absorbing aerosol trends at middle and high

latitudes may be affected by industrial activity. If so, they

would contribute to long-term changes in UV amount.

The current TOMS UV algorithm accounts for absorbing

and scattering aerosol effects, and could be used for

extended trend estimates by combining the Nimbus-7/-

TOMS and Earth-Probe/TOMS data.

5.4.3 Estimates from Radiative Transfer

Models

Sabziparvar et al. (1999) have calculated the climatology of surface UV radiation using the discreteordinate RT model of Stamnes et al. (1988) modified by

Forster (1995). Calculations were made for daily doses

of UV-B (280-315 nm) and UV-A (315-400 nm), and

weighted with various biological action spectra.

Calculations were performed for the four midseason

months: January, April, July, and October. Clear-sky and

aerosol-free irradiance values were calculated using

TOMS total ozone averages as input. All-sky conditions

were calculated using average cloud distributions measured by the International Satellite Cloud Climatology

Project (Rossow and Schiffer, 1991) and aerosol optical

depth of 0.1 at 400 nm. Figure 5-7 shows the annual

average of daily dose calculations for erythemally

weighted UV irradiance.

The same RT model was used by Sabziparvar et al.

(1998) to calculate changes in surface UV radiation due

to changes in ozone (both stratospheric and tropospheric)

and tropospheric aerosols since preindustrial times (1850).

Decreases in erythemally weighted irradiance of up to 9%

were calculated for tropical regions as a result of increases

in tropospheric ozone. Decreases in erythemally weighted

irradiance of 4-8% from increased tropospheric ozone

were calculated north of 20° in the Northern Hemisphere

during July. Increases of erythemal irradiance of 70% at

southern high latitudes in October and 10% at northern

high latitudes were calculated as a result of stratospheric

ozone depletion. Effects of increases in absorbing

aerosols and sulfate aerosols were also considered, and it

was concluded that an increase in aerosols has decreased

erythemally weighted UV by up to 7% locally and 2% on

a global average. The study does not include long-term

changes in cloud cover or surface albedo.

5.4.4 Estimates from Ground-Based

Ancillary Data

Gantner et al. (2000) reconstructed noontime irradiance values at 300, 305, 310, and 320 nm for clear-sky

and no-snow conditions at Hohenpeissenberg going back

to 1968. Statistical dependencies of irradiance on total

ozone were determined from measurements taken
SURFACE ULTRAVIOLET RADIATION

5.24

between 1990 and 1997 when spectral UV, total ozone,

and global and diffuse radiation data are available, and

the relationships were used to extrapolate estimates of UV

irradiance backward in time. For all months between

March and September, statistically significant long-term

increases in clear-sky noontime values of UV were found

that increase with decreasing wavelengths.

The studies by Kaurola et al. (2000) (using data

from Norrkoping, Sweden; Jokioinen, Finland; and Belsk,

Poland) and den Outer et al. (2000) (using data from

Bilthoven, Netherlands) took another approach to reconstruct past UV data that includes effects of clouds and

other scattering processes. RT models were used to calculate clear-sky daily integral values of erythemally

weighted UV irradiance and global solar irradiance (280-

4000 nm) using total ozone, other meteorological data,

and the daily variation of SZA as input. The ratio of the

daily integral of pyranometer measurements to the modeled clear-sky integral was used to quantify scattering

processes that are related statistically to variations in UV

observed in recent years. The relationships were used to

estimate daily UV values under all weather conditions

backward in time to 1979 using pyranometer data and

satellite total ozone data, or to the 1960s if ground-based

total ozone measurements were available. In general, the

reconstructed datasets show that erythemally weighted

UV has increased since the 1970s and that the increases

and year-to-year temporal variations are consistent with

satellite estimates for the same sites (Kaurola et al., 2000).

Diaz et al. (2000, 2002) demonstrated a method

that calculates the daily integral of UV radiation (at specific wavelength intervals, broadband or biologically

weighted irradiance) from daily average total ozone, daily

average SZA (for daylight hours), and the daily integral

of broadband measurements (UV-B, biologically

weighted UV, or global solar pyranometer). A multilinear

Figure 5-7. Climatology of annual average erythemal dose (in KJoule/(m2 day)). Values are computed from

total ozone and cloud satellite data. Adapted from Sabziparvar et al. (1999).
SURFACE ULTRAVIOLET RADIATION

5.25

regression model was used to fit the daily UV values as a

function of the three measured independent variables, and

regression coefficients were determined using data from

South Pole, Antarctica; Ushuaia, Argentina; and San

Diego, U.S. It was shown that monthly averages of UV

irradiance between 303 and 308 nm can be predicted to

an accuracy of 4-5% using broadband radiation data. The

method was applied to TOMS total ozone data and pyranometer data from South Pole and Barrow, Alaska. The

regression was determined using data from 1993 to 1998,

and results of the regression were used to estimate UV

radiation (303-308 nm) back to October 1979. Significant

increases (~300%) were found for the month of October

at South Pole and lower values (between 90% and 120%)

for April at Barrow between 1979 and 1996.

Results of statistical models described by Fioletov

et al. (1997, 2001) and McArthur et al. (1999) establish

empirical dependencies of UV radiation (both at individual wavelengths and spectral intervals) on SZA, total

ozone, global solar radiation, dewpoint temperature, and

snow cover. The relationships were determined from

observations at six sites in Canada (Edmonton, Winnipeg,

Churchill, Toronto, Montreal, and Halifax) where total

ozone, spectral UV irradiance, and global solar irradiance

(pyranometer) measurements have been made since the

early 1990s. It was shown that monthly mean values of

erythemal irradiance for summer months (May-August)

can be estimated from the ancillary data with a standard

error (1-sigma) of 4.2%. The statistical relationships have

been used to estimate hourly UV irradiance values dating

back to the mid 1960s at Toronto, Churchill, and

Edmonton when both total ozone and pyranometer data

are available (Fioletov et al., 2001). The hourly values

were integrated to make daily total and noontime (within

1 hour of noon) values, and monthly averages of the daily

and noontime values were determined. Monthly average

daily UV doses for summer months (May, June, July, and

August) predicted from the ancillary data agree with

simultaneous spectral measurements with an rms difference of 3.3%. This agreement is better than the 4% rms

difference for summer monthly ground-based spectral

measurements and satellite estimates (Fioletov et al.,

2002). Trends in UV for individual wavelengths and

weighted spectral intervals were determined for the period

from 1979 to 1997 and are shown in Figure 5-8. It is of

interest to note that trends in the daily integrated values

and noontime values are essentially the same. The annual

trend values at wavelengths of 310 nm or less and for the

erythemally weighted UV are all statistically significant

to the 2-sigma level.

In addition to the estimation of past hour-by-hour

UV irradiance, the ancillary data can be used to quantify

and distinguish between trends in UV that are caused by

factors other than long-term changes in total ozone.

Figure 5-8 shows that Churchill had statistically significant trends at all wavelengths, including those with

insignificant ozone absorption (325 nm). The positive

trend of about 7% per decade at 325 nm is due to the combined effect of an increase of days with snow cover

(causing an increase in UV irradiance by about 2% per

decade) and a decrease in hours of cloudiness (causing

an increase of about 5% per decade) that occurred at

Churchill over the period (1979-1997). The 11% per

decade increase of erythemally weighted irradiance was

found to consist of +2% per decade from increasing snow

cover, +5% per decade from decreasing clouds, and +4%

per decade from decreasing ozone.

A continuous UV time series made up of irradiance

measurements from recent years (1990s) and of irradiance estimates using ancillary data from earlier years (as

early as the mid-1960s) is useful to quantify extreme

events that are pertinent to some biological studies.

Fioletov et al. (2001) demonstrated that the incidence of

extreme events has increased since 1970 by showing that

the number of hours per year the UV Index is above a

threshold value has increased at Toronto, Edmonton, and

Churchill (Figure 5-9). This is in accordance with the

findings for an alpine site (Seckmeyer et al., 1997), which

concluded the same with the support of the ancillary measurements made at Hohenpeissenberg.

5.5 EXPECTATIONS OF UV IN THE FUTURE

Just as the climatology of surface UV radiation has

changed in the past, there will likely be changes over the

next several decades. The future evolution of UV irradiance will depend on changes in total ozone as well as other

atmospheric variables that are influenced by changes in

climate. Quantification of predicted changes in total

ozone, influenced both by halogen chemistry as well as

other climate variables, is discussed and presented in previous chapters in this Assessment. Here we address the

changes in UV irradiance that are likely to occur over the

next 50 years, both as a result of the expected recovery of

the ozone layer (Section 5.5.1) and variations driven by

other variables influenced by climate change that directly

affect surface UV irradiance (Section 5.5.2).

5.5.1 Links with the Recovery of the Ozone

Layer

It is well known that a reduction in total ozone

column leads to increased levels of surface UV irradiance

if other factors remain constant. This relationship has
SURFACE ULTRAVIOLET RADIATION

5.26

been clearly demonstrated and reported in previous

Assessments (WMO, 1992, 1995, 1999).

There has been a significant amount of work done

to predict the future of the ozone layer based on considerations of the reduction of atmospheric chlorine and

bromine compounds as well as the interactions with climate change. Previous chapters in this Assessment discuss these future ozone scenarios in detail. Future UV

levels as a result of changes in ozone have been predicted

(Taalas et al., 2000b; Reuder et al., 2001), but, because

surface UV radiation is dependent on the many complex

scattering and absorbing processes discussed previously

in this chapter, prediction of future UV radiation is more

difficult and leads to deviating results about the probable

UV levels over the coming decades.

Chapter 4 of this Assessment presents results of

several two-dimensional (2-D) models that simulate the

decline of total ozone between 1979 and 2000 and project

its recovery using several halogen and climate change

scenarios between 2000 and 2050. The derived ozone

projections do not include polar effects (e.g., increased

occurrence of polar stratospheric clouds) and their influence at higher latitudes. The results have been used with

RT model calculations to estimate future UV irradiance

over the next 50 years. Monthly values of total ozone at

5° latitude intervals were averaged from the output of

seven 2-D model runs for the MA2 scenario (described in

Chapter 4) and used as input to the RT model used by

NASA to estimate surface UV irradiance from TOMS

satellite data (Herman et al., 1996, 1999; Krotkov et al.,

1998, 2001). Monthly values for noontime erythemal irradiance were determined using the total ozone from

Chapter 4, and noontime SZA at the middle of each month,

for clear skies and low (3%) albedo. Figure 5-10 shows

 May-August 

1 0

2 0

3 0

4 0

Trend (% per decade)

Churchill (59 N, 94 W )

Year

Churchill (59 N, 94 W )

1 0

2 0

3 0

4 0

Trend (% per decade)

Edmonton (54 N, 114 W ) Edmonton (54 N, 114 W)

300 305 310 315 320 325

1 0

2 0

3 0

4 0

Trend (% per decade)

CIE DUV

Wavelen th nm

Toronto (44 N, 79 W)

300305310315320325 CIE DUV

Toronto (44 N, 79 W)

Wavelength (nm)

Figure 5-8. Trends in UV irradiance at 300,

305, 310, 315, 320, and 325 nm, and trends

in erythemally (CIE) and DNA weighted

irradiance for 1979-1997, for May-August

and year-round, at Churchill (58.75°N), Edmonton (53.55°N), and Toronto (43.75°N),

with 2s error estimates indicated. Solid

squares are trends in daily integrals, and

open squares are trends from values within

±1 hour of solar noon. Adapted from

Fioletov et al. (2001).
the departure (in percent) of these clear-sky calculations

of noontime erythemally weighted UV from the 1980

mean values for the period from 1979 to 2050 for the

months of January, April, July, and October.

Results of the 2-D model calculations (Figure 5-

10) indicate that generally there will be small changes in

UV radiation over the next decade. However, in the longer

term (next 50 years) UV levels will return to values similar to those that were present in 1980, provided that all

other factors (clouds, snow cover, aerosols, etc.) are

assumed to remain constant. It should be stressed that this

assumption is an unlikely scenario. Past changes of surface UV that are associated with changes other than total

ozone have occurred at some sites, as illustrated in Figure

5-5 (Thessaloniki) and Figure 5-8 (Churchill). It should

further be emphasized that there have been many surprises

in the occurrence of extreme events in the past due to the

complex and nonlinear behavior of the atmosphere. These

surprises cannot be included in such a projection.

Impacts of greenhouse gases and halogenated

species on future UV levels have been studied by Taalas

et al. (2000b). The UV scenarios are based on ozone predictions from 3-D chemistry-climate calculations carried

out by three different models that were discussed in

Chapter 3: the Goddard Institute for Space Studies (GISS;

Shindell et al., 1998) model, the U.K. Meteorological

Office (Austin et al., 2000) model, and the German

Aerospace Institute model.

Figure 5-11 shows results of future erythemal UV

levels relative to 1979-1992 averages calculated using the

GISS model ozone predictions for April in the Northern

Hemisphere and October in the Southern Hemisphere

(from Taalas et al., 2000b). For the Northern Hemisphere

in April there are enhancements up to 90% at 70°N calculated for the 2010-2020 time period that reduce to <25%

in 2040-2050. For the Southern Hemisphere in October

the enhancements are up to 100% over Antarctica for the

2010-2020 period and reduce to <50% by 2040-2050.

Maximum increases in the annual (not shown) northern

high-latitude UV doses were estimated to be up to 14% in

2010-2020, and reduce to 2% in 2040-2050 because of

the expected recovery of the ozone layer. At southern high

latitudes, 40% annual maximum enhancements are

expected during 2010-2020 that reduce to 27% by 2040-

2050.

The following points should be noted regarding the

rather large deviations stated above, particularly for the

Arctic. First, the above values (Figure 5-11) are relative

to the period 1979-1992. Since the models indicate that

we are nearing minimum ozone values in 2002 (Chapter

3), deviations relative to today’s values are appreciably

less than those stated above for the 2010-2020 period and

are negative for the 2040-2050 period. Second, the deviations are spatially dependent, and the values stated above

are at places where the maxima are predicted to occur.

When UV levels are averaged from 60°N to 90°N, the

GISS model agrees reasonably well with other models

(Taalas et al., 2000b). Finally, the GISS prediction of minimum UV levels for the Northern Hemisphere would be

higher than those of other 3-D models, since the GISS

SURFACE ULTRAVIOLET RADIATION

5.27

0

50

100

150

200

250 Number of Hours

UV Index>6

UV Index>7

Toronto (44ON, 79OW)

0

50

100

150

200

250 Number of Hours

UV Index>5

UV Index>6

Edmonton (54ON, 114OW)

1965 1970 1975 1980 1985 1990 1995

0

20

40

60

80

100 Number of Hours

UV Index>5

UV Index>6

Churchill (59ON, 94OW)

Figure 5-9. The number of hours per year that the

UV Index exceeded 6 and 7 at Toronto, 5 and 6 at

Edmonton, and 5 and 6 at Churchill. These numbers are determined from reconstructed erythemal

UV irradiance determined from total ozone and other

ancillary data. Adapted from Fioletov et al. (2001).
SURFACE ULTRAVIOLET RADIATION

model predicts the lowest ozone levels for the Arctic

(Chapter 3).

5.5.2 Factors Related to Climate Change

The effects of climate change on surface UV radiation are twofold. The first (indirect) effects are caused by

changes in climate that affect total ozone. The consequential changes in total ozone are addressed in Chapters 3 and

4 of this Assessment, and the impact of these changes on

surface UV radiation are discussed in Section 5.5.1 above.

The second effects of climate change on UV are direct

effects (e.g., changes in clouds, aerosols, snow cover, etc.)

and those are discussed in this section.

The recent IPCC report (IPCC, 2001) lists many

factors that are likely to change due the increased emission of greenhouse gases. Many of these factors have

already started to change or are likely to have changed in

the past. In addition to these feedbacks that are already

discussed in previous chapters (e.g., the higher likelihood

of polar stratospheric clouds and their associated role in

enhancing the chemical destruction of ozone), there are

other factors related to climate change that are directly

influencing UV irradiance at the Earth’s surface.

In Section 5.2.1.5 it is discussed that snow albedo

leads to a significant enhancement of UV irradiance above

the surface due to multiple scattering. The IPCC report

(IPCC, 2001) states that the sea ice cover in the Northern

Hemisphere has reduced significantly and is likely to

reduce more in the future. In addition, the snow cover

and the extent of glaciers in mountainous areas have been

decreasing. Consequently, it can be predicted that UV

irradiance above the surface will decrease in these areas

for seasons where snow cover or sea ice is reduced.

Although UV irradiance above the surface is enhanced by

snow-covered areas, a decrease in sea ice and snow cover

5.28

Figure 5-10. Calculations of the departures (in percent) from 1980 UV noontime clear-sky irradiance levels

between 1979 and 2050 for the months of (a) January, (b) April, (c) July, and (d) October. The average of

seven model runs calculating total ozone from the MA2 scenario discussed in Chapter 4 were used as input to

the TOMS radiative transfer model (Herman et al., 1996, 1999; Krotkov et al., 1998, 2001). These results

show that UV levels will have insignificant variation over the next 10 years and will return to nearly the same

levels as those present in 1980 by 2050. It should be stressed that these calculations assume all variables

other than total ozone remain constant, which is unlikely.
SURFACE ULTRAVIOLET RADIATION

5.29

will result in an increase of UV dose for organisms living

under water and on land areas previously covered by snow.

According to the IPCC report (2001), the largest reductions of snow albedo will be at high latitudes in spring and

summer, and consequently surface UV irradiance will

likely be reduced by the decrease in albedo at high latitudes in spring and summer in some areas of the Northern

Hemisphere.

Figure 5-11. Results of erythemal UV changes due to ozone changes predicted by the Goddard Institute for

Space Studies (GISS) 3-D chemistry-climate model for April in the Northern Hemisphere and October in the

Southern Hemisphere. Future increases of greenhouse gases are expected to delay the recovery of the

ozone layer (Chapter 3), and the resulting enhanced UV values at high latitudes are predicted to persist for

about two more decades. The indicated changes are relative to the time period 1979-1992, so the changes

relative to present-day levels are less than those shown for 2010-2020 and would likely be negative for 2040-

2050. Adapted from Taalas et al. (2000b).
SURFACE ULTRAVIOLET RADIATION

5.30

Another statement of the IPCC 2001 report has an

impact on surface UV: it is thought that cloudiness has

increased by about 2% over the past century and is likely

to continue increasing in the future. The effects of clouds

on UV irradiance depend both on their amount and type.

The observed increase of cirrus clouds in the Northern

Hemisphere is likely to reduce UV irradiance slightly.

However, this increase would lead mainly to a redistribution of UV radiance from the direct beam to diffuse. More

important is the likely increase of overall global cloudiness over the next century. The effect of this increase will

lead to a reduction of surface UV globally. However, it

should be noted that it is very likely that cloudiness will

increase in some regions and decrease in others. Recent

results by Chen et al. (2002) and Wielicki et al. (2002)

demonstrate unexplained decreases in cloud cover in the

tropics. These changes will have a great influence on the

UV irradiance locally, possibly exceeding those caused

by changes in ozone.

Weather extremes may occur more frequently than

they have in the past century. In the future, climate

episodes of high precipitation might alternate with

extremely dry episodes in many parts of the Earth. Higher

UV irradiance can be expected during dry spells than

during wet spells, and this may be problematic for some

biological species. At present it is not possible to predict

these changes quantitatively. Nevertheless the biological

impact of these changes might exceed those caused by the

expected changes in stratospheric ozone.

Aerosol effects on surface UV radiation may be

larger than previously thought, and may affect large areas

of the globe because of the following:

(1) Many anthropogenic aerosols absorb in the UV

(Jacobson, 1999).

(2) Altitude gradients of UV irradiance are generally

larger than expected (Lenoble et al., 2002; McKenzie

et al., 2001b; Seckmeyer et al., 1997). The relatively

high altitude gradients could be partly explained by

the attenuation of UV irradiance by aerosols near the

ground.

(3) Satellite estimates of surface UV-B irradiance are too

large over large areas of the globe, possibly caused

by a widespread influence of the aerosols that is not

accounted for in the satellite estimates (McKenzie et

al., 2001a; Fioletov et al., 2002).

Over the first half of the 20th century, aerosol

extinction increased (Liu et al., 1991; Sabziparvar et al.,

1998). In recent years aerosol extinction has started to

decrease with a cleaner troposphere in some areas (e.g.,

Krzys ́cin and Puchalski, 1998; Zerefos et al., 1998). What

might happen in the future is open to conjecture,

Table 5-1. Factors related to climate change that influence UV irradiance at the surface globally.

Regional changes are expected to be over a much larger range than the numbers given here. Also the relative importance of the factors may be different in different regions (e.g., snow cover would become a major

factor in mountainous areas). It should be emphasized that there have been many surprises in the occurrence of extreme events in the past due to the complex and nonlinear behavior of the atmosphere. These surprises cannot be included in such a projection.

Correlation with UV

Change

Predicted Change in Globally Averaged

Erythemal UV from Present

2010 Long Term

Major Factors

Stratospheric ozone Negative 0 to –3% Figure 5-10

Cloudiness Negative ±1% Negative

Aerosols Negative ±3% Negative

Tropospheric ozone Negative ±1% –2 to +7%

Minor Factors

Aviation (IPCC, 1999) ±1% ±1%

Ozone profile ±1% ±1%

Stratospheric temperature Negative Positive Positive

Snow cover and sea ice Positive Negative in some areas and seasons

Tropospheric temperature Negative Negative Negative

Weather extremes Positive peak values
SURFACE ULTRAVIOLET RADIATION

5.31

depending on the disparate forcings of increased population, increased awareness of environmental issues, and

reduced reliance on fossil fuels. However, it is now likely

that UV radiation at the ground will increase in the coming

years, before the expected recovery of the ozone layer

might take place.

The IPCC report on impacts of aviation on the

atmosphere (IPCC, 1999) considered changes in UV that

could result from increased aviation activity both in the

upper troposphere and lower stratosphere. These changes

were generally found to be less than ±1% due to effects

that influence UV negatively or positively.

Table 5-1 shows a summary of factors related to

climate change that influence UV irradiance at the surface on a global scale. An attempt has been made to quantify changes that may occur both in the shorter term (next

10 years) and longer term (to 2050). It should be emphasized that the UV changes are expected to be larger in

certain areas under specific situations (e.g., by changing

pollution, cloud cover, or surface albedo).

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SURFACE ULTRAVIOLET RADIATION

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SURFACE ULTRAVIOLET RADIATION

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Appendix 5A

SPECTRAL DATA AVAILABLE FROM DATABASES

5A.1 World Ozone and Ultraviolet Radiation Data Centre (WOUDC) a

Station Name Instrument Country Institute b Lat. (°N) Long. (°E) Start

Date

7 Kagoshima Brewer Japan JMA 31.5 130.5 1/01/91

12 Sapporo Brewer Japan JMA 43.02 141.3 1/01/91

14 Tateno Brewer Japan JMA 36.02 140.07 1/01/90

18 Alert Brewer Canada MSC 82.47 ñ62.35 6/09/95

21 Edmonton (Stony Pl.) Brewer Canada MSC 53.52 ñ114.13 3/19/92

24 Resolute Brewer Canada MSC 74.69 ñ95.01 3/14/91

31 Mauna Loa Brewer (double) United States MSC 19.5 ñ155.6 1/01/98

31 Mauna Loa Brewer United States MSC 19.5 ñ155.6 3/24/97

31 Mauna Loa Bentham United States NIWA 19.5 ñ155.6 11/01/97

65 Toronto Brewer Canada MSC 43.75 ñ79.5 3/01/89

76 Goose Bay Brewer Canada MSC 53.15 ñ60.51 1/01/97

77 Churchill Brewer Canada MSC 58.72 ñ94.1 3/30/92

95 Taipei Brewer Taiwan CWBT 24.99 121.49 3/06/92

101 Syowa Brewer Japan JMA ñ69.03 ñ39.55 1/01/93

111 Amundsen-Scott Biospherical c Antarctica NSF ñ89.96 ñ24.8 1/01/91

190 Naha Brewer Japan JMA 26.17 127.65 1/01/91

199 Barrow Biospherical c United States NSF 71.32 ñ156.6 1/01/91

239 San Diego Biospherical c United States NSF 32.45 ñ117.11 1/01/92

241 Saskatoon Brewer Canada MSC 52.08 ñ106.74 1/01/91

256 Lauder Bentham New Zealand NIWA ñ45.03 169.68 1/01/94

261 Thessaloniki Brewer Greece AUTH 40.52 22.97 10/01/89

268 Arrival Heights Biospherical c Antarctica NSF ñ77.83 166.67 1/01/89

290 Saturna Island Brewer Canada MSC 48.75 ñ123.16 10/03/90

292 Palmer Biospherical c Antarctica NSF ñ64.75 ñ64.05 1/01/91

301 Ispra Brewer (double) Italy JRC-EC 45.8 8.63 1/01/01

306 Chengkung Brewer Taiwan CWBT 23.07 121.34 1/02/92

307 Obninsk Brewer Russia IEM-SPA 55.09 35.97 5/06/93

319 Montreal (Dorval) Brewer Canada MSC 45.45 ñ73.78 3/06/93

320 Winnipeg Brewer Canada MSC 49.87 ñ97.27 7/06/92

321 Halifax (Bedford) Brewer Canada MSC 44.65 ñ63.67 7/02/92

331 Poprad-Ganovce Brewer Slovakia SHMI 49 20.29 1/01/94

332 Pohang Brewer Korea KMA 36 129.35 1/27/94

338 Bratts Lake (Regina) Brewer Canada MSC 50.18 ñ104.74 1/01/95

339 Ushuaia Biospherical c Argentina NSF ñ54.85 ñ68.31 1/01/90

353 Reading Optronic United Kingdom UMIST 51.42 0.96 1/27/94

a Website: http://www.msc-smc.ec.gc.ca/woudc/

b JMA, Japan Meteorological Agency; MSC, Meteorological Service of Canada; NIWA, National Institute of Water and Atmospheric Research;

CWBT, Central Weather Bureau of Taiwan; NSF, National Science Foundation; AUTH, Aristotle University of Thessaloniki; JRC-EC, Joint Research

Centre-European Commission; IEM-SPA, Institute of Experimental Meteorology-Scientific Production Association; SHMI, Slovak

Hydrometeorological Institute; KMA, Korea Meteorological Administration; UMIST, University of Manchester Institute of Science and Technology.

c Data from Biospherical Instruments are also available via the website http://www.biospherical.com/nsf
SURFACE ULTRAVIOLET RADIATION

5.44

5A.2 European Ultraviolet Database (EUVDB) a

No. Station Name Instrument Country Institute b Lat. (°N) Long. (°E) Start

Date

1 Sonnblick Bentham Austria BOKU 47.05 12.97 Feb. 96

2 Vienna Bentham Austria BOKU 48.23 16.35 Mar. 98

3 Uccle, Brussels Jobin & Yvon Belgium IASB 50.80 4.36 Mar. 93

4 Jokioinen Brewer Finland FMI 60.81 23.50 Apr. 95

5 Sodankyl‰ Brewer Finland FMI 67.37 26.63 Apr. 90

6 Villar St. Pancrace Jobin & Yvon France USTL 44.90 6.65 Sep. 99

7 Villeneuve díAsq Jobin & Yvon France USTL 50.37 3.01 Apr. 97

8 Neuherberg Bentham Germany BfS 48.22 11.58 Jan. 01

9 Offenbach Bentham Germany BfS 50.10 8.75 Jul. 01

10 Hohenpeissenberg Brewer Germany DWD 47.80 11.02 Jan. 95

11 Lindenberg Brewer Germany DWD 52.22 14.12 Jan. 95

12 Potsdam Brewer Germany DWD 52.36 13.08 Jan. 95

13 Garmisch-Partenkirchen Bentham Germany IFU 47.48 11.07 Apr. 94

14 Zugspitze Bentham Germany IFU 7.42 10.98 Apr. 94

15 Thessaloniki Brewer Greece LAP 40.52 22.97 Oct. 89

16 Lampedusa Italy ENEA 35.50 12.60 Feb. 98

17 Ispra Brewer Italy JRC 45.81 8.63 Jan. 92

18 Rome Brewer Italy URO 41.90 12.52 Feb. 92

19 And ̄ya Bentham Norway NILU 69.48 16.02 Feb. 98

20 Trondheim Optronic Norway NTNU 64.43 10.47 Mar. 97

21 Troms ̄ Brewer Norway UT 69.66 18.93 May. 94

22 Belsk Brewer Poland IGFPAS 51.83 20.78 Jan. 93

23 Azores Brewer Portugal MI 38.66 ñ27.22

24 Funchal (Madeira Is.) Brewer Portugal MI 32.64 ñ16.89 Jan. 97

25 Lisbon Brewer Portugal MI 38.77 ñ9.15 Jun. 00

26 Penhas Douradas Brewer Portugal MI 40.42 ñ7.55

27 Izana Brewer Spain INM 28.49 ñ16.50 Jan. 95

28 Norrkoping Brewer Sweden SMHI 58.58 16.15 Sep. 91

29 Vindeln Brewer Sweden SMHI 64.23 19.77 Jul. 96

30 De Bilt Brewer Netherlands KNMI 52.10 5.18 Jan. 94

31 LSO (Bilthoven) Dilcon Netherlands RIVM 52.12 5.20 Jan. 96

32 Reading Optronics United Kingdom UMIST 51.45 ñ0.93 Jan. 94

Only permanent stations are listed; campaign data are not included.

Data from Neuherberg, Offenbach, Lampedusa, Andoya, Azores, and Penhas Douradas will be submitted in the coming months.

a Website: http://www.muk.uni-hannover.de/EDUCE

b BOKU, Universit‰t f ̧r Bodenkultur Wein; IASB, Institut díAÈronomie Spatiale de Belgique; FMI, Finnish Meteorological Institute; USTL,

UniversitÈ des Sciences et Technologies de Lille; BfS, Bundesamt f ̧r Strahlenschutz; DWD, Deutscher Wetterdienst; IFU, Atmosph‰rische

Umweltforschung; LAP, Laboratory of Atmospheric Physics; ENEA, Ente per le Nuove Technologie líEnergia e líAmbiente; JRC, Joint Research

Centre, European Commission; URO, University of Rome; NILU, Norsk Institutt for Luftforskning; NTNU, Norges Teknisk-Naturvitenskapelige

Universitet; UT, University of Troms ̄; IGFPAS, Institute of Geophysics, Polish Academy of Sciences; MI, Institute of Meteorology, Portugal; INM,

Instituto Nacional de Meteorologia; SMHI, Swedish Meteorological and Hydrological Institute; KNMI, Koninklijk Nederlands Meteorologische

Instituut; RIVM, Rijkinstituut voor Volksgezandheid en Milieu; UMIST, University of Manchester Institute of Science and Technology.
SURFACE ULTRAVIOLET RADIATION

5.45

Appendix 5B

INTERNET ADDRESSES FOR UV SITES

5B.1 General UV Information

WMO UV radiation page http://titan.srrb.noaa.gov/UV/

TOMS (Ozone and UV satellite data) http://toms.gsfc.nasa.gov/

NSF UV monitoring network, U.S. http://www.biospherical.com/NSF/

USDA UV-B monitoring and research program http://uvb.nrel.colostate.edu/

NIWA, New Zealand http://www.niwa.co.nz/services/uvozone/

EPA, U.S. http://www.epa.gov/uvnet/ and

http://oz.physast.uga.edu/

EPA SunWise School Program http://www.epa.gov/sunwise/

WMO Ozone Bulletins http://www.wmo.ch/web/arep/ozone.html

Center for International Earth Science http://sedac.ciesin.columbia.edu/ozone/

Information Network (CIESIN), U.S.

Atmospheric Sciences Research Center, U.S. http://uvb.asrc.cestm.albany.edu/

World Radiation Centre http://wrdc-mgo.nrel.gov/

5B.2 International UV Projects

World Ozone and UV Data Center http://www.msc-smc.ec.gc.ca/woudc/

European Cooperation in the field of Scientific http://159.213.57.69/uvweb

and Technical research (COST-713)

SUVDAMA, European Commission http://www.ozone.fmi.fi/SUVDAMA/

EDUCE, European Commission http://www.muk.uni-hannover.de/EDUCE/

Thematic Network For Ultraviolet Measurements http://metrology.hut.fi/uvnet/

QASUME, European Communities http://lap.physics.auth.gr/qasume/

5B.3 Radiative Transfer Codes

Tropospheric Ultraviolet Visible (TUV) http://www.acd.ucar.edu/TUV/

library for Radiative Transfer (libRadtran) http://www.libradtran.org/

System for Transfer of Atmospheric Radiation http://www.meteo.physik.uni-muenchen.de/strahlung/uvrad/

(STAR) Star/STARinfo.htm

Fast and Easy Radiative Transfer (FASTRT) http://zardoz.nilu.no/~olaeng/fastrt/fastrt.html

Santa Barbara DISTORT Atmospheric http://arm.mrcsb.com/sbdart/

Radiative Transfer (SBDART)

5B.4 Extraterrestrial Spectra

Solar Ultraviolet Spectral Irradiance Monitor http://wwwsolar.nrl.navy.mil/susim_atlas_data.html

(SUSIM)

Solar Stellar Irradiance Comparison Experiment http://lasp.colorado.edu/solstice/ and 

(SOLSTICE) http://www-uars.gsfc.nasa.gov/cdrom_main.html

McMath/Pierce at Kitt Peak ftp://ftp.noao.edu/fts/fluxatl/
SURFACE ULTRAVIOLET RADIATION

5.46

5B.5 General UV Index Information

ìUV-Index for the public,î COST-713 http://www.bag.admin.ch/strahlen/nonionisant/pdf/e/dossier-cost.pdf

EPA, U.S. http://www.epa.gov/sunwise/uvindex.html

NOAA/EPA, U.S. http://www.cpc.ncep.noaa.gov/products/stratosphere/

uv_index/index.html

University of Vienna, Austria http://www-med-physik.vu-wien.ac.at/uv/uv_online.htm

The UV Index Sun Awareness Program, Canada http://www.msc-smc.ec.gc.ca/uvindex/index_e.html

World Health Organization (WHO) http://www.who.int/inf-fs/en/fact133.html

Intersun/WHO http://www.who.int/peh-uv/

5B.6 Internet Sites with UV Information by Country

Argentina http://www.conae.gov.ar/caratula.html and

http://www.meteofa.mil.ar

Australia (BOM) http://www.bom.gov.au/info/about_uvb.shtml

Australia (ARPANSA) http://www.arpansa.gov.au/is_uvindex.htm

Australia (SunSmart) http://www.sunsmart.com.au/

Austria (University Vienna) http://www-med-physik.vu-wien.ac.at/uv/uv_online.htm

Austria (University Innsbruck) http://www.uibk.ac.at/projects/uv-index/

Canada http://www.ec.gc.ca/ozone/

Czech Republic http://www.chmi.cz/meteo/ozon/o3uvb-e.html

European Commission (JRC) http://ecuv.jrc.it/

Finland http://www.fmi.fi/research_atmosphere/atmosphere_2.html

France http://www.securite-solaire.org/

Germany (DWD) http://www.uv-index.de/

Germany (BfS) http://www.bfs.de/uvi/index.htm

Greece http://lap.physics.auth.gr/uvindex/

Italy http://www.lamma.rete.toscana.it/previ/eng/uvhtm/uvnew0.html

Japan http://www.shiseido.co.jp/e/e9708uvi/html/

Mexico http://www.sima.com.mx/t1msn_valle_de_mexico/uv-index.asp

New Zealand http://www.niwa.co.nz/services/uvozone/

Norway http://uvnett.nrpa.no/uv/

Poland http://www.imgw.pl/

Portugal http://www.meteo.pt/uv/uvindex.htm

Spain http://infomet.am.ub.es/uv_i_ozo/uvi.html

Sweden http://www.smhi.se/weather/uvindex/sv/uvprog.htm

Switzerland http://www.bag.admin.ch/strahlen/nonionisant/uv/d/index.php

United States http://www.cpc.ncep.noaa.gov/products/stratosphere/uv_index/
TWENTY QUESTIONS AND ANSWERS

ABOUT THE OZONE LAYER

Lead Author:

D.W. Fahey

A draft of this component of the Assessment was reviewed and discussed by the 74 scientists who attended the Panel

Review Meeting for the 2002 ozone assessment (Les Diablerets, Switzerland, 24-28 June 2002). In addition, subsequent

contributions, reviews, or comments were provided by the following individuals: A.-L.N. Ajavon, D.L. Albritton, S.O.

Andersen, P.J. Aucamp, G. Bernhard, M.P. Chipperfield, J.S. Daniel, S.B. Diaz, E.S. Dutton, C.A. Ennis, P.J. Fraser, R.-S.

Gao, R.R. Garcia, M.A. Geller, S. Godin-Beekmann, M. Graber, J.B. Kerr, M.K.W. Ko, M.J. Kurylo, G.L. Manney, K.

Mauersberger, M. McFarland, G. Mégie, S.A. Montzka, R. Müller, E.R. Nash, P.A. Newman, S.J. Oltmans, M.

Oppenheimer, L.R. Poole, G. Poulet, M.H. Proffitt, W.J. Randel, A.R. Ravishankara, C.E. Reeves, R.J. Salawitch, M.L.

Santee, G. Seckmeyer, D.J. Siedel, K.P. Shine, C.C. Sweet, A.F. Tuck, G.J.M. Velders, R.T. Watson, and R.J. Zander. 
TWENTY QUESTIONS AND ANSWERS ABOUT THE OZONE LAYER

Contents

page

INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.1

I. OZONE IN OUR ATMOSPHERE

Q1. What is ozone and where is it in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.3

Q2. How is ozone formed in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.4

Q3. Why do we care about atmospheric ozone? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.5

Q4. Is total ozone uniform over the globe? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.6

Q5. How is ozone measured in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.7

II. THE OZONE DEPLETION PROCESS

Q6. What are the principal steps in stratospheric ozone depletion caused by

human activities? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.8

Q7. What emissions from human activities lead to ozone depletion? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.10

Q8. What are the reactive halogen gases that destroy stratospheric ozone? . . . . . . . . . . . . . . . . . . . . . . . . . Q.13

Q9. What are the chlorine and bromine reactions that destroy stratospheric ozone? . . . . . . . . . . . . . . . . . . Q.16

Q10. Why has an “ozone hole” appeared over Antarctica when ozone-depleting gases are present

throughout the stratosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.18

III. STRATOSPHERIC OZONE DEPLETION

Q11. How severe is the depletion of the Antarctic ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.21

Q12. Is there depletion of the Arctic ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.23

Q13. How large is the depletion of the global ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.25

Q14. Do changes in the Sun and volcanic eruptions affect the ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . Q.26

IV. CONTROLLING OZONE-DEPLETING GASES

Q15. Are there regulations on production of ozone-depleting gases? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.28

Q16. Has the Montreal Protocol been successful in reducing ozone-depleting gases

in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.30

V. IMPLICATIONS OF OZONE DEPLETION

Q17. Does depletion of the ozone layer increase ground-level ultraviolet radiation? . . . . . . . . . . . . . . . . . . Q.32

Q18. Is depletion of the ozone layer the principal cause of climate change? . . . . . . . . . . . . . . . . . . . . . . . . . Q.34

VI. STRATOSPHERIC OZONE IN THE FUTURE

Q19. How will recovery of the ozone layer be detected? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.36

Q20. When is the ozone layer expected to recover? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.37

ADDITIONAL TOPICS

• Understanding Stratospheric Ozone Depletion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.9

• Heavier-Than-Air CFCs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.12

• Global Ozone Dobson Network . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.15

• The Discovery of the Antarctic Ozone Hole . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.20

• Replacing the Loss of “Good” Ozone in the Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.24
TWENTY QUESTIONS

Q.1

Ozone is a very small part of our atmosphere, but its

presence is nevertheless vital to human well-being.

Most ozone resides in the upper part of the atmosphere. This region, called the stratosphere, is more than

10 kilometers (6 miles) above Earth’s surface. There,

about 90% of atmospheric ozone is contained in the

“ozone layer,” which shields us from harmful ultraviolet

light from the Sun.

However, it was discovered in the mid-1970s that

some human-produced chemicals could destroy ozone and

deplete the ozone layer. The resulting increase in ultraviolet radiation at Earth’s surface can increase the incidences

of skin cancer and eye cataracts.

Following the discovery of this environmental issue,

researchers focused on a better understanding of this threat

to the ozone layer. Monitoring stations showed that the

abundances of the ozone-depleting chemicals were

steadily increasing in the atmosphere. These trends were

linked to growing production and use of chemicals like

chlorofluorocarbons (CFCs) for refrigeration and air

conditioning, foam blowing, and industrial cleaning.

Measurements in the laboratory and the atmosphere characterized the chemical reactions that were involved in

ozone destruction. Computer models employing this

information could then predict how much ozone depletion was occurring and how much more could occur in the

future.

Observations of the ozone layer itself showed that

depletion was indeed occurring. The most severe and

most surprising ozone loss was discovered to be over

Antarctica. It is commonly called the “ozone hole”

because the ozone depletion is so large and localized. A

thinning of the ozone layer also has been observed over

other regions of the globe, such as the Arctic and northern

middle latitudes.

The work of many scientists throughout the world

has provided a basis for building a broad and solid scientific understanding of the ozone depletion process. With

this understanding, we know that ozone depletion is occurring and why. And, most important, it has become clear

that, if ozone-depleting gases were to continue to accumulate in the atmosphere, the result would be more ozone

layer depletion.

In response to the prospect of increasing ozone depletion, the governments of the world crafted the 1987 United

Nations Montreal Protocol as a global means to address a

global issue. As a result of the broad compliance with the

Protocol and its Amendments and Adjustments and,

importantly, the industrial development of more “ozonefriendly” substitutes for the now-controlled chemicals,

the total global accumulation of ozone-depleting gases

has slowed and begun to decrease. This has reduced the

risk of further ozone depletion. Now, with continued compliance, we expect recovery of the ozone layer in the late

21st century. International Day for the Preservation of the

Ozone Layer, 16 September, is now celebrated on the day

the Montreal Protocol was agreed upon.

This is a story of notable achievements: discovery,

understanding, decisions, and actions. It is a story written

by many: scientists, technologists, economists, legal

experts, and policymakers. And, dialogue has been a key

ingredient.

To help foster a continued interaction, this component of the Scientific Assessment of Ozone Depletion:

2002 presents 20 questions and answers about the oftencomplex science of ozone depletion. The questions

address the nature of atmospheric ozone, the chemicals

that cause ozone depletion, how global and polar ozone

depletion occur, and what could lie ahead for the ozone

layer. A brief answer to each question is first given in

italics; an expanded answer then follows. The answers

are based on the information presented in the 2002 and

earlier Assessment reports. These reports and the answers

provided here were all prepared and reviewed by a large

international group of scientists. 1

INTRODUCTION

1 A draft of this component of the Assessment was reviewed and discussed by the 74 scientists who attended the Panel Review Meeting for the 2002

ozone assessment (Les Diablerets, Switzerland, 24-28 June 2002). In addition, subsequent contributions, reviews, or comments were provided by the

following individuals: A.-L.N. Ajavon, D.L. Albritton, S.O. Andersen, P.J. Aucamp, G. Bernhard, M.P. Chipperfield, J.S. Daniel, S.B. Diaz, E.S.

Dutton, C.A. Ennis, P.J. Fraser, R.-S. Gao, R.R. Garcia, M.A. Geller, S. Godin-Beekmann, M. Graber, J.B. Kerr, M.K.W. Ko, M.J. Kurylo, M.

McFarland, G.L. Manney, K. Mauersberger, G. Mégie, S.A. Montzka, R. Müller, E.R. Nash, P.A. Newman, S.J. Oltmans, M. Oppenheimer, L.R.

Poole, G. Poulet, M.H. Proffitt, W.J. Randel, A.R. Ravishankara, C.E. Reeves, R.J. Salawitch, M.L. Santee, G. Seckmeyer, D.J. Siedel, K.P. Shine,

C.C. Sweet, A.F. Tuck, G.J.M. Velders, R.T. Watson, and R.J. Zander. 
TWENTY QUESTIONS

Q.3

Ozone is a gas that is naturally present in our atmosphere. Because an ozone molecule contains three oxygen

atoms (see Figure Q1-1), it has a chemical formula of O3.

Ozone was discovered in laboratory experiments in the

mid-1800s. Ozone’s presence in the atmosphere was later

discovered using chemical and optical measurement

methods. The word ozone is derived from the Greek word

ozein, meaning “to smell.” Ozone has a pungent odor that

allows ozone to be detected even in very low amounts.

Ozone will rapidly react with many chemical compounds

and is explosive in concentrated amounts. Electrical discharges are generally used to make ozone for industrial

processes including air and water purification and

bleaching of textiles and food products.

Ozone location. Most ozone (about 90%) is found

in the stratosphere, a region that begins about 10-16 kilometers (6-10 miles) above Earth’s surface and extends

up to about 50 kilometers (31 miles) altitude (see Figure

Q1-2). The stratosphere begins at higher altitudes (16

kilometers) in the tropics than in the polar regions (10

kilometers). Most ozone resides in the stratosphere in

what is commonly known as the “ozone layer.” The

remaining ozone, about 10%, is found in the troposphere,

which is the lowest region of the atmosphere between

Earth’s surface and the stratosphere. 

Ozone abundance. Ozone molecules have a relatively low abundance in the atmosphere. In the stratosphere near the peak of the ozone layer, there are up to

12,000 ozone molecules for every billion air molecules

(1 billion = 1000 million). Most air molecules are either

oxygen (O2) or nitrogen (N2) molecules. In the troposphere near Earth’s surface, ozone is even less abundant,

with a typical range of 20 to 100 ozone molecules for

each billion air molecules. The highest surface values

are a result of ozone formed in air polluted by human

activities. 

As an illustration of the low relative abundance of

ozone in our atmosphere, one can consider bringing all the

ozone molecules in the troposphere and stratosphere down

to Earth’s surface and uniformly distributing these molecules into a gas layer over the globe. The resulting layer

of pure ozone would have a thickness of less than onehalf centimeter (about one-quarter inch). Altitude (kilometers)

Ozone concentration

Altitude (miles)

0

5

10

15

20

25

30

35

Stratospheric

Ozone

Tropospheric

Ozone

20

15

10

5

Ozone Layer

Ozone

increases

from pollution

Ozone in the Atmosphere

Figure Q1-2. Atmospheric ozone. Ozone is present

throughout the lower atmosphere. Most ozone resides

in the stratospheric “ozone layer” above Earth’s surface.

Increases in ozone occur near the surface as a result of

pollution from human activities. 

Oxygen

Atom (O)

Ozone

Molecule (O3)

Oxygen

Molecule (O2 )

Ozone and Oxygen

Figure Q1-1. Ozone and oxygen. A molecule of ozone

(O3) contains three oxygen (O) atoms bound together.

Oxygen molecules (O2), which constitute 21% of Earth’s

atmosphere, contain two oxygen atoms bound together.

I. OZONE IN OUR ATMOSPHERE

Q1: What is ozone and where is it in the atmosphere? 

Ozone is a gas that is naturally present in our atmosphere. Each ozone molecule contains three atoms of oxygen and

is denoted chemically as O3. Ozone is found primarily in two regions of the atmosphere. About 10% of atmospheric

ozone is in the troposphere, the region closest to Earth (from the surface to about 10-16 kilometers (6-10 miles)). The

remaining ozone (90%) resides in the stratosphere, primarily between the top of the troposphere and about 50 kilometers (31 miles) altitude. The large amount of ozone in the stratosphere is often referred to as the “ozone layer.”
TWENTY QUESTIONS

Q.4

Stratospheric ozone. Stratospheric ozone is naturally formed in chemical reactions involving ultraviolet

sunlight and oxygen molecules, which make up 21% of

the atmosphere. In the first step, sunlight breaks apart one

oxygen molecule (O2) to produce two oxygen atoms

(2 O) (see Figure Q2-1). In the second step, each atom

combines with an oxygen molecule to produce an ozone

molecule (O3). These reactions occur continually wherever ultraviolet sunlight is present in the stratosphere. As

a result, the greatest ozone production occurs in the tropical stratosphere. 

The production of stratospheric ozone is balanced by

its destruction in chemical reactions. Ozone reacts continually with a wide variety of natural and humanproduced chemicals in the stratosphere. In each reaction,

an ozone molecule is lost and other chemical compounds

are produced. Important reactive gases that destroy ozone

are those containing chlorine and bromine (see Q8). 

Some stratospheric ozone is transported down into

the troposphere and can influence ozone amounts at

Earth’s surface, particularly in remote unpolluted regions

of the globe.

Tropospheric ozone. Near Earth’s surface, ozone is

produced in chemical reactions involving naturally occurring gases and gases from pollution sources. Production

reactions primarily involve hydrocarbon and nitrogen

oxide gases and require sunlight. Fossil fuel combustion

is a primary pollution source for tropospheric ozone production. The surface production of ozone does not significantly contribute to the abundance of stratospheric ozone.

The amount of surface ozone is too small and the transport of surface air to the stratosphere is not effective

enough. As in the stratosphere, ozone in the troposphere

is destroyed in naturally occurring chemical reactions and

in reactions involving human-produced chemicals.

Tropospheric ozone can also be destroyed when ozone

reacts with a variety of surfaces such as those of soils and

plants.

Balance of chemical processes. Ozone abundances

in the stratosphere and troposphere are determined by the

balance between chemical processes that produce and

destroy ozone. The balance is determined by the amounts

of reacting gases and by how the rate or effectiveness of

the various reactions varies with sunlight intensity, location in the atmosphere, temperature, and other factors. As

atmospheric conditions change to favor ozone-production

reactions in a certain location, ozone abundances will

increase. Similarly, if conditions change to favor reactions that destroy ozone, its abundances will decrease.

The balance of production and loss reactions combined

with atmospheric air motions determines the global distribution of ozone on time scales of days to many months.

Global ozone has decreased in the last decades because

the amounts of reactive gases containing chlorine and

bromine have increased in the stratosphere (see Q13).

Overall reaction: 3O2 sunlight 2O3

Stratospheric Ozone Production

Ultraviolet

Sunlight

Step

1 +

Step

2

Figure Q2-1. Stratospheric ozone production. Ozone

is naturally produced in the stratosphere in a two-step

process. In the first step, ultraviolet sunlight breaks apart

an oxygen molecule to form two separate oxygen atoms.

In the second step, these atoms then undergo a binding

collision with other oxygen molecules to form two ozone

molecules. In the overall process, three oxygen molecules react to form two ozone molecules.

Q2: How is ozone formed in the atmosphere?

Ozone is formed throughout the atmosphere in multistep chemical processes that require sunlight. In the stratosphere, the process begins with the breaking apart of an oxygen molecule (O2 ) by ultraviolet radiation from

the Sun. In the lower atmosphere (troposphere), ozone is formed in a different set of chemical reactions involving hydrocarbons and nitrogen-containing gases.
TWENTY QUESTIONS

Q.5

All ozone molecules are chemically identical, with

each containing three oxygen atoms. However, ozone in

the stratosphere has very different environmental consequences for humans and other life forms than ozone in the

troposphere near Earth’s surface.

Good ozone. Stratospheric ozone is considered

“good” for humans and other life forms because it absorbs

ultraviolet (UV)-B radiation from the Sun (see Figure Q3-

1). If not absorbed, UV-B would reach Earth’s surface in

amounts that are harmful to a variety of life forms. In

humans, as their exposure to UV-B increases, so does their

risk of skin cancer (see Q17), cataracts, and a suppressed

immune system. The UV-B exposure before adulthood

and cumulative exposure are both important factors in the

risk. Excessive UV-B exposure also can damage terrestrial plant life, single-cell organisms, and aquatic ecosystems. Other UV radiation, UV-A, which is not absorbed

significantly by ozone, causes premature aging of the skin.

The absorption of UV-B radiation by ozone is a source

of heat in the stratosphere. This helps to maintain the stratosphere as a stable region of the atmosphere with temperatures

increasing with altitude. As a result, ozone plays a key role in

controlling the temperature structure of Earth’s atmosphere.

Protecting good ozone. In the mid-1970s, it was discovered that some human-produced gases could cause stratospheric ozone depletion (see Q6). Ozone depletion increases

harmful UV-B amounts at Earth’s surface. Global efforts have

been undertaken to protect the ozone layer through the regulation of ozone-depleting gases (see Q15 and Q16).

Bad ozone. Ozone is also formed near Earth’s surface

in natural chemical reactions and in reactions caused by the

presence of human-made pollutant gases. Ozone produced

by pollutants is “bad” because more ozone comes in direct

contact with humans, plants, and animals. Increased levels

of ozone are generally harmful to living systems because

ozone reacts strongly to destroy or alter many other molecules. Excessive ozone exposure reduces crop yields and

forest growth. In humans, ozone exposure can reduce lung

capacity; cause chest pains, throat irritation, and coughing;

and worsen pre-existing health conditions related to the

heart and lungs. In addition, increases in tropospheric

ozone lead to a warming of Earth’s surface (see Q18). The

negative effects of increasing tropospheric ozone contrast

sharply with the positive effects of stratospheric ozone as

an absorber of harmful UV-B radiation from the Sun.

Reducing bad ozone. Reducing the emission of pollutants can reduce “bad” ozone in the air surrounding

humans, plants, and animals. Major sources of pollutants

include large cities where fossil fuel consumption and industrial activities are greatest. Many programs around the globe

have already been successful in reducing the emission of

pollutants that cause near-surface ozone production.

Natural ozone. Ozone is a natural component of the

clean atmosphere. In the absence of human activities on

Earth’s surface, ozone would still be present near the surface

and throughout the troposphere and stratosphere. Ozone’s

chemical role in the atmosphere includes helping to remove

other gases, both those occurring naturally and those emitted

by human activities. If all the ozone were to be removed from

the lower atmosphere, other gases such as methane, carbon

monoxide, and nitrogen oxides would increase in abundance.

Ozone Layer

Earth

UV Protection by the Ozone Layer

Figure Q3-1. UV-B protection by the ozone layer.

The ozone layer resides in the stratosphere and surrounds the entire Earth. UV-B radiation (280- to 315-

nanometer (nm) wavelength) from the Sun is partially

absorbed in this layer. As a result, the amount reaching

Earth’s surface is greatly reduced. UV-A (315- to 400-

nm wavelength) and other solar radiation are not

strongly absorbed by the ozone layer. Human exposure to UV-B increases the risk of skin cancer,

cataracts, and a suppressed immune system. UV-B

exposure can also damage terrestrial plant life, singlecell organisms, and aquatic ecosystems.

Q3: Why do we care about atmospheric ozone?

Ozone in the stratosphere absorbs some of the Sun’s biologically harmful ultraviolet radiation. Because of this

beneficial role, stratospheric ozone is considered “good ozone.” In contrast, ozone at Earth’s surface that is

formed from pollutants is considered “bad ozone” because it can be harmful to humans and plant and animal

life. Some ozone occurs naturally in the lower atmosphere where it is beneficial because ozone helps remove

pollutants from the atmosphere.
TWENTY QUESTIONS

Q.6

Total ozone. Total ozone at any location on the

globe is found by measuring all the ozone in the atmosphere directly above that location. Total ozone includes

that present in the stratospheric ozone layer and that present throughout the troposphere (see Figure Q1-2). The

contribution from the troposphere is generally only about

10% of total ozone. Total ozone values are often reported in Dobson units, denoted “DU.” Typical values vary

between 200 and 500 DU over the globe (see Figure Q4-

1). A total ozone value of 500 DU, for example, is equivalent to a layer of pure ozone gas on Earth’s surface having a thickness of only 0.5 centimeters (0.2 inches). 

Global distribution. Total ozone varies strongly

with latitude over the globe, with the largest values occurring at middle and high latitudes (see Figure Q4-1). This

is a result of winds that circulate air in the stratosphere,

moving tropical air rich in ozone toward the poles in fall

and winter. Regions of low total ozone occur at polar latitudes in winter and spring as a result of the chemical

destruction of ozone by chlorine and bromine gases (see

Q11 and Q12). The smallest values of total ozone (other

than in the Antarctic in spring) occur in the tropics in all

seasons, in part because the thickness of the ozone layer

is smallest in the tropics.

Natural variations. The variations of total ozone

with latitude and longitude come about for two reasons.

First, natural air motions mix air between regions of the

stratosphere that have high ozone values and those that

have low ozone values. Air motions also increase the vertical thickness of the ozone layer near the poles, which

increases the value of total ozone in those regions.

Tropospheric weather systems can temporarily reduce the

thickness of the stratospheric ozone layer in a region,

lowering total ozone at the same time. Second, variations

occur as a result of changes in the balance of chemical

production and loss processes as air moves to new locations over the globe. Reductions in solar ultraviolet radiation exposure, for example, will reduce the production of

ozone.

Scientists have a good understanding of how chemistry and air motion work together to cause the observed

large-scale features in total ozone such as those seen in

Figure Q4-1. Ozone changes are carefully monitored by

a large group of investigators using satellite, airborne, and

ground-based instruments. The analysis of these observations helps scientists to estimate the contribution of

human activities to ozone depletion.

22 June 1999

Global Satellite Maps of Total Ozone

22 December 1999

200 250 300 350 400 450 500

Total Ozone (Dobson units)

Q4: Is total ozone uniform over the globe?

No, the total amount of ozone above the surface of Earth varies with location on time scales that range from

daily to seasonal. The variations are caused by stratospheric winds and the chemical production and destruction of ozone. Total ozone is generally lowest at the equator and highest near the poles because of the seasonal wind patterns in the stratosphere.

Figure Q4-1. Total ozone. A total ozone value is

obtained by measuring all the ozone that resides in the

atmosphere over a given location on Earth’s surface.

Total ozone values shown here are reported in

“Dobson units” as measured by a satellite instrument

from space. Total ozone varies with latitude, longitude,

and season, with the largest values at high latitudes

and the lowest values in tropical regions. Total ozone

at most locations varies with time on a daily to seasonal basis as ozone-rich air is moved about the globe by

stratospheric winds. Low total ozone values over

Antarctica in the 22 December image represent the

remainder of the “ozone hole” from the 1999 Antarctic

winter/spring season (see Q11).
TWENTY QUESTIONS

Q.7

The abundance of ozone in the atmosphere is measured by a variety of techniques (see Figure Q5-1). The

techniques make use of ozone’s unique optical and chemical properties. There are two principal categories of

measurement techniques: direct and remote. Ozone

measurements by these techniques have been essential in

monitoring changes in the ozone layer and in developing

our understanding of the processes that control ozone

abundances.

Direct measurements. Direct measurements of

atmospheric ozone abundance are those that require air to

be drawn directly into an instrument. Once inside an

instrument, ozone can be measured by its absorption of

ultraviolet (UV) light or by the electrical current produced

in an ozone chemical reaction. The latter approach is used

in the construction of “ozonesondes,” which are lightweight ozone-measuring modules suitable for launching

on small balloons. Small balloons ascend far enough in

the atmosphere to measure ozone in the stratospheric

ozone layer. Ozonesondes are launched weekly at many

locations around the world. Direct ozone-measuring

instruments using optical or chemical detection schemes

are also used routinely on board research aircraft to

measure the distribution of ozone in the troposphere and

lower stratosphere. High-altitude research aircraft can

reach the ozone layer at most locations over the globe and

can reach farthest into the layer at high latitudes in polar

regions. Ozone measurements are also being made on

some commercial aircraft. 

Remote measurements. Remote measurements of

ozone abundance are obtained by detecting the presence

of ozone at large distances away from the instrument.

Most remote measurements of ozone rely on its unique

absorption of UV radiation. Sources of UV radiation that

can be used are the Sun and lasers. For example, satellites use the absorption of UV sunlight by the atmosphere

or the absorption of sunlight scattered from the surface

of Earth to measure ozone over the globe on a near-daily

basis. A network of ground-based detectors measures

ozone by the amount of the Sun’s UV light that reaches

Earth’s surface. Other instruments measure ozone using

its absorption of infrared or visible radiation or its emission of microwave or infrared radiation. Total ozone

amounts and the altitude distribution of ozone can be

obtained with remote measurement techniques. Lasers

are routinely deployed at ground sites or on board aircraft

to detect ozone over a distance of many kilometers along

the laser light path.

Large

aircraft

Measuring Ozone in the Atmosphere

Satellites

Ground-based

systems

Laser

beams

Balloon

sondes

High-altitude

aircraft

Figure Q5-1. Ozone measurements. Ozone is measured throughout the atmosphere with instruments on the

ground and on board aircraft, high-altitude balloons, and

satellites. Some instruments measure ozone directly in

sampled air and others measure ozone remotely some

distance away from the instrument. Instruments use

optical techniques with the Sun and lasers as light

sources or use chemical reactions that are unique to

ozone. Measurements at many locations over the globe

are made weekly to monitor total ozone amounts.

Q5: How is ozone measured in the atmosphere?

The amount of ozone in the atmosphere is measured by instruments on the ground and carried aloft in balloons,

aircraft, and satellites. Some measurements involve drawing air into an instrument that contains a system for

detecting ozone. Other measurements are based on ozone’s unique absorption of light in the atmosphere. In

that case, sunlight or laser light is carefully measured after passing through a portion of the atmosphere containing ozone.
TWENTY QUESTIONS

Q.8

Emission, accumulation, and transport. The principal steps in stratospheric ozone depletion caused by

human activities are shown in Figure Q6-1. The process

begins with the emission of halogen source gases at

Earth’s surface (see Q7). Halogen source gases include

manufactured gases containing chlorine or bromine that

are released to the atmosphere by a variety of human

activities. Chlorofluorocarbons (CFCs), for example, are

important chlorine-containing gases. These source gases

accumulate in the lower atmosphere (troposphere) and

are eventually transported to the stratosphere. The accumulation occurs because most source gases are unreactive

in the lower atmosphere (troposphere). Small amounts of

these gases dissolve in ocean waters. 

Some emissions of halogen gases come from natural

sources (see Q7). These emissions also accumulate in the

troposphere and are transported to the stratosphere. 

Conversion, reaction, and removal. Halogen

source gases do not react directly with ozone. Once in the

stratosphere, halogen source gases are chemically converted to reactive halogen gases by ultraviolet radiation

from the Sun (see Q8). These reactive gases chemically

destroy ozone in the stratosphere (see Q9). The average

depletion of total ozone attributed to reactive gases is estimated to be small in the tropics and up to about 10% at

middle latitudes (see Q13). In polar regions, the presence

of polar stratospheric clouds greatly increases the abundance of the most reactive halogen gases (see Q10). This

results in ozone destruction in polar regions in winter and

II. THE OZONE DEPLETION PROCESS

Q6: What are the principal steps in stratospheric ozone depletion caused by human activities?

The initial step in the depletion of stratospheric ozone by human activities is the emission of ozone-depleting

gases containing chlorine and bromine at Earth’s surface. Most of these gases accumulate in the lower atmosphere because they are unreactive and do not dissolve readily in rain or snow. Eventually, the emitted gases

are transported to the stratosphere where they are converted to more reactive gases containing chlorine and

bromine. These more reactive gases then participate in reactions that destroy ozone. Finally, when air returns

to the lower atmosphere, these reactive chlorine and bromine gases are removed from Earth’s atmosphere by

rain and snow.

Figure Q6-1. Principal steps in stratospheric ozone

depletion. The stratospheric ozone depletion process

begins with the emission of halogen source gases at

Earth’s surface and ends when reactive halogen gases

are removed by rain and snow in the troposphere and

deposited on Earth’s surface. In the stratosphere, the

reactive halogen gases, namely chlorine monoxide

(ClO) and bromine monoxide (BrO), destroy ozone.

Principal Steps in the Depletion

of Stratospheric Ozone

Halogen source gases are emitted at Earth's

surface by human activities and natural processes.

Halogen source gases accumulate in the

atmosphere and are distributed throughout the

lower atmosphere by winds and other air motions.

Halogen source gases are transported to the

stratosphere by air motions.

Most halogen source gases are converted in the

stratosphere to reactive halogen gases in chemical

reactions involving ultraviolet radiation from the Sun.

Air containing reactive halogen gases returns

to the troposphere and these gases are removed

from the air by moisture in clouds and rain.

Reactive halogen gases cause chemical

depletion of stratospheric total ozone over the

globe except at tropical latitudes.

Emissions

Accumulation

Transport

Chemical reaction

Conversion

Removal

Polar stratospheric clouds increase ozone

depletion by reactive halogen gases,

causing severe ozone loss in polar regions

in winter and spring.
TWENTY QUESTIONS

Q.9

spring (see Q11 and Q12). After a few years, air in the

stratosphere returns to the troposphere, bringing along

reactive halogen gases. These gases are then removed

from the atmosphere by rain and other precipitation and

deposited on Earth’s surface. This removal brings to an

end the destruction of ozone by chlorine and bromine

atoms that were first released to the atmosphere as components of halogen source gas molecules.

Tropospheric conversion. Halogen source gases

with short lifetimes (see Q7) undergo significant chemical conversion in the troposphere, producing reactive

halogen gases and other compounds. Source gas molecules that are not converted accumulate in the troposphere

and are transported to the stratosphere. Because of

removal by precipitation, only small portions of the reactive halogen gases produced in the troposphere are also

transported to the stratosphere. Important examples of

gases that undergo some tropospheric removal are the

HCFCs, which are used as substitute gases for other halogen source gases (see Q15 and Q16), bromoform, and

gases containing iodine (see Q7).

Understanding Stratospheric Ozone Depletion

Scientists learn about ozone destruction through a combination of laboratory studies, computer

models, and stratospheric observations. In laboratory studies scientists are able to discover and evaluate

individual chemical reactions that also occur in the stratosphere. Chemical reactions between two gases

follow well-defined physical rules. Some of these reactions occur on the surfaces of particles formed in the

stratosphere. Reactions have been studied that involve a wide variety of molecules containing chlorine,

bromine, fluorine, and iodine and other atmospheric constituents such as oxygen, nitrogen, and hydrogen.

These studies show that there exist several reactions involving chlorine and bromine that can directly or

indirectly cause ozone destruction in the atmosphere. 

With computer models, scientists can examine the overall effect of a large group of known reactions

under the chemical and physical conditions found in the stratosphere. These models include winds, air

temperatures, and the daily and seasonal changes in sunlight. With such analyses, scientists have shown

that chlorine and bromine can react in catalytic cycles in which one chlorine or bromine atom can destroy

many ozone molecules. Scientists use model results to compare with past observations as a test of our

understanding of the atmosphere and to evaluate the importance of new reactions found in the laboratory.

Computer models also enable scientists to explore the future by changing atmospheric conditions and other

model parameters. 

Scientists make stratospheric observations to find out what gases are present in the stratosphere and

at what concentrations. Observations have shown that halogen source gases and reactive halogen gases

are present in the stratosphere at expected amounts. Ozone and chlorine monoxide (ClO), for example,

have been observed extensively with a variety of instruments. Instruments on the ground and on board

satellites, balloons, and aircraft detect ozone and ClO at a distance using optical and microwave signals.

High-altitude aircraft and balloon instruments detect both gases directly in the stratosphere. For example,

these observations show that ClO is present at elevated amounts in the Antarctic and Arctic stratospheres

in the winter/spring season, when the most severe ozone depletion occurs.
TWENTY QUESTIONS

Q.10

Human-produced chlorine and bromine gases.

Human activities cause the emission of halogen source

gases that contain chlorine and bromine atoms. These

emissions into the atmosphere ultimately lead to stratospheric ozone depletion. The source gases that contain only

carbon, chlorine, and fluorine are called “chlorofluorocarbons,” usually abbreviated as CFCs. CFCs, along with carbon tetrachloride (CCl4) and methyl chloroform (CH3CCl3),

are the most important chlorine-containing gases that are

emitted by human activities and destroy stratospheric ozone

Q7: What emissions from human activities lead to ozone depletion?

Certain industrial processes and consumer products result in the atmospheric emission of “halogen source

gases.” These gases contain chlorine and bromine atoms, which are known to be harmful to the ozone layer.

For example, the chlorofluorocarbons (CFCs) and hydrochlorofluorocarbons (HCFCs), once used in almost all

refrigeration and air conditioning systems, eventually reach the stratosphere where they are broken apart to

release ozone-depleting chlorine atoms. Other examples of human-produced ozone-depleting gases are the

“halons,” which are used in fire extinguishers and which contain ozone-depleting bromine atoms. The production and consumption of all principal halogen source gases by human activities are regulated worldwide under

the Montreal Protocol.

Total chlorine amount (parts per trillion)

Chlorine source gases

Primary Sources of Chlorine and Bromine

 for the Stratosphere in 1999

16%

32%

23%

12%

7%

5%

1%

4%

0

3400

3000

2000

1000

Bromine source gases

Total bromine amount (parts per trillion)

0

20

15

10

5

(CH3CCl3)

(e.g., HCFC-22 = CHClF2)

(CCl2FCClF2)

15%

27-42%

5-20%

20%

14%

4%

Human-made sources Natural sources Human-made sources

Natural

sources

Other gases

Methyl chloroform

HCFCs

CFC-113

Carbon tetrachloride (CCl4)

CFC-11 (CCl3F)

CFC-12 (CCl2F2)

Methyl chloride (CH3Cl)

Methyl bromide (CH3Br)

Halon-1211 (CBrCIF2)

Halon-1301 (CBrF3)

Other halons

Very-short lived gases

 (e.g., bromoform = CHBr3)

Figure Q7-1. Stratospheric source gases. A variety

of gases transport chlorine and bromine into the stratosphere. These gases, called halogen source gases, are

emitted from natural sources and human activities. For

chlorine, human activities account for most that reaches

the stratosphere. The CFCs are the most abundant of

the chlorine-containing gases released in human activities. Methyl chloride is the most important natural

source of chlorine. For bromine that reaches the stratosphere, halons and methyl bromide are the largest

sources. Both gases are released in human activities.

Methyl bromide has an additional natural source.

Natural sources are a larger fraction of the total for

bromine than for chlorine. HCFCs, which are substitute

gases for CFCs and also are regulated under the

Montreal Protocol, are a small but growing fraction of

chlorine-containing gases. The abundance scale shows

that the amount of chlorine in the stratosphere is greater

than that of bromine by about 170 times. (The unit “parts

per trillion” is a measure of the relative abundance of

a gas: 1 part per trillion indicates the presence of one

molecule of a gas per trillion other molecules.)
TWENTY QUESTIONS

Q.11

(see Figure Q7-1). Chlorine-containing gases have been

used in many applications including refrigeration, air conditioning, foam blowing, aerosol propellants, and cleaning

of metals and electronic components. These activities have

typically caused the emission of halogen-containing gases

to the atmosphere. 

Another category of halogen source gases contains

bromine. The most important of these are the “halons”

and methyl bromide (CH3Br). Halons are halogenated

hydrocarbon gases originally developed to extinguish

fires. Halons are widely used to protect large computers,

military hardware, and commercial aircraft engines.

Because of these uses, halons are often directly released

into the atmosphere. Halon-1211 and Halon-1301 are the

most abundant halons emitted by human activities (see

Figure Q7-1). Methyl bromide, used primarily as an

agricultural fumigant, is also a significant source of

bromine to the atmosphere.

Human emissions of chlorine- and bromine-containing

gases have increased substantially since the middle of the

20th century (see Q16). The result has been global ozone

depletion with the greatest losses occurring in polar regions

(see Q11 to Q13).

Natural sources of chlorine and bromine. There

are two halogen source gases present in the stratosphere

that have large natural sources. These are methyl chloride

(CH3Cl) and methyl bromide (CH3Br), both of which are

emitted by oceanic and terrestrial ecosystems. Natural

sources of these two gases contribute about 16% of the

chlorine currently in the stratosphere, and about 27-42%

of the bromine (see Figure Q7-1). Very short-lived gases

containing bromine, such as bromoform (CHBr3), are also

released to the atmosphere by the oceans. The estimated

contribution of these gases to stratospheric bromine

(about 15%) is uncertain at this time. Changes in the natural sources of chlorine and bromine since the middle of

the 20th century are not the cause of observed ozone

depletion. 

Lifetimes and emissions. After emission, halogen

source gases are either removed from the atmosphere or

undergo chemical conversion. The time to remove or convert about 60% of a gas is often called its atmospheric

“lifetime”. Lifetimes vary from less than 1 year to 100

years for the principal chlorine- and bromine-containing

gases (see Table Q7-1). Gases with the shortest lifetimes

(e.g., the HCFCs, methyl bromide, methyl chloride, and

the very short-lived gases) are significantly destroyed in

the troposphere, and therefore only a fraction of the emitted gas contributes to ozone depletion in the stratosphere. 

The amount of a halogen source gas present in the

atmosphere depends on the lifetime of the gas and the

amount emitted to the atmosphere. Emissions vary greatly for the principal source gases, as indicated in Table

Q7-1. Emissions of most gases regulated by the

Montreal Protocol have decreased since 1990, and emissions from all regulated gases are expected to decrease in

the coming decades (see Q16).

Ozone Depletion Potential. The halogen source

gases in Figure Q7-1 are also known as “ozone-depleting

substances” because they are converted in the stratosphere to reactive gases containing chlorine and bromine

(see Q8). Some of these reactive gases participate in

reactions that destroy ozone (see Q9). Ozone-depleting

substances are compared in their effectiveness to destroy

stratospheric ozone using the “Ozone Depletion

Potential” (ODP), as listed in Table Q7-1. A gas with a

larger ODP has a greater potential to destroy ozone over

its lifetime in the atmosphere. The ODP is calculated on

a “per mass” basis for each gas relative to CFC-11, which

has an ODP defined to be 1. Halon-1211 and Halon-1301

have ODPs significantly larger than CFC-11 and most

other emitted gases. This results because bromine is

much more effective overall (about 45 times) on a peratom basis than chlorine in chemical reactions that

destroy ozone in the stratosphere. The gases with small

Table Q7-1. Atmospheric lifetimes, emissions,

and Ozone Depletion Potentials of halogen

source gases. a

Halogen Lifetime Global Ozone

Source (years) Emissions Depletion

Gas in 2000 Potential

(gigagrams (ODP)

per year) b

Chlorine

CFC-12 100 130-160 1

CFC-113 85 10-25 1

CFC-11 45 70-110 1

Carbon 26 70-90 0.73

tetrachloride

HCFCs 1-26 340-370 0.02-0.12

Methyl chloroform 5 ~20 0.12

Methyl chloride 1.3 3000-4000 0.02

Bromine

Halon-1301 65 ~3 12

Halon-1211 16 ~10 6

Methyl bromide 0.7 160-200 0.38

Very short-lived Less c c

gases than 1

a Includes both human activities and natural sources. 

b 1 gigagram = 109 grams = 1000 metric tons. 

c No reliable estimate available.
TWENTY QUESTIONS

Q.12

ODP values generally have short atmospheric lifetimes.

The production and consumption of all principal halogen

source gases by humans are regulated under the provisions of the Montreal Protocol (see Q15).

Fluorine and iodine. Fluorine and iodine are also

halogen atoms. Most of the source gases in Figure Q7-1

also contain fluorine atoms in addition to chlorine or

bromine. After the source gases undergo conversion in

the stratosphere (see Q6), the fluorine content of these

gases is left in chemical forms that do not destroy ozone.

Iodine is a component of several gases that are naturally

emitted from the oceans. Although iodine will participate

in ozone destruction reactions, these iodine-containing

gases are largely removed in the troposphere by natural

processes before the gases can reach the stratosphere. 

Other gases. Other gases that influence stratospheric ozone abundances also have increased in the

stratosphere as a result of human activities. Important

examples are methane (CH4) and nitrous oxide (N2O),

which react to form water vapor and reactive hydrogen

and nitrogen oxides, respectively, in the stratosphere.

These reactive products also participate in the production

and loss balance of stratospheric ozone (see Q2). The

overall affect of these other gases on ozone is much

smaller than that caused by increases in chlorine- and

bromine-containing gases from human activities.

Heavier-Than-Air CFCs

CFCs and other halogen source gases reach the stratosphere despite the fact they are “heavier than

air.” All the principal source gases are emitted and accumulate in the lower atmosphere (troposphere). Air

containing the emitted halogen gases is in continual motion as a result of winds and convection. Air motions

ensure that the source gases are horizontally and vertically well mixed throughout the troposphere in a

matter of months. It is this well-mixed air that enters the lower stratosphere from upward air motions in

tropical regions, bringing with it source gas molecules emitted from a wide variety of locations on Earth’s

surface.

Atmospheric measurements confirm that halogen source gases with long atmospheric lifetimes are

well mixed in the troposphere and are present in the stratosphere (see Figure Q8-2). The amounts found in

these regions are consistent with the emissions estimates reported by industry and government.

Measurements also show that gases that are “lighter than air,” such as hydrogen (H2) and methane (CH4),

are also well mixed in the troposphere, as expected. Only at altitudes well above the troposphere and

stratosphere (above 85 kilometers (53 miles)), where much less air is present, do heavy gases begin to

separate from lighter gases as a result of gravity.
TWENTY QUESTIONS

Q.13

Reactive gases containing the halogens chlorine and

bromine lead to the chemical destruction of stratospheric

ozone. Halogen-containing gases present in the stratosphere can be divided into two groups: halogen source

gases and reactive halogen gases. The source gases are

emitted at Earth’s surface by natural processes and by

human activities (see Q7). Once they reach the stratosphere, the halogen source gases chemically convert to

form the reactive halogen gases.

Reactive halogen gases. The chemical conversion

of halogen source gases, which involves ultraviolet sunlight and other chemical reactions, produces a number of

reactive halogen gases. These reactive gases contain all

of the chlorine and bromine atoms originally present in

the source gases.

The most important reactive chlorine- and brominecontaining gases that form in the stratosphere are shown

in Figure Q8-1. Away from polar regions, the most abundant are hydrogen chloride (HCl) and chlorine nitrate

(ClONO2). These two gases are considered reservoir

gases because they do not react directly with ozone but

can be converted to the most reactive forms that do chemically destroy ozone. The most reactive forms are chlorine monoxide (ClO) and bromine monoxide (BrO), and

chlorine and bromine atoms (Cl and Br). A large fraction

of available stratospheric bromine is generally in the form

of BrO whereas usually only a small fraction of stratospheric chlorine is in the form of ClO. In polar regions,

the reservoirs ClONO2 and HCl undergo a further conversion on polar stratospheric clouds (see Q10) to form ClO.

In that case, ClO becomes a large fraction of available

reactive chlorine.

Reactive chlorine observations. Reactive chlorine

gases have been observed extensively in the stratosphere

with both direct and remote measurement techniques. The

measurements from space at middle latitudes displayed in

Figure Q8-2 are representative of how chlorine-containing

gases change between the surface and the upper stratosphere. Available chlorine (see red line in Figure Q8-2) is

the sum of chlorine contained in halogen source gases and

CFC-12

Methyl

chloroform

Very short-lived

bromine gases

Carbon

tetrachloride

Methyl bromide

Methyl

chloride

CFC-113

CFC-11 Halon-1211

Stratospheric Halogen Gases

Chemical Hydrogen bromide (HBr)

Bromine nitrate (BrONO2)

Hydrogen chloride (HCl)

Chlorine nitrate (ClONO2)

Chlorine monoxide (ClO)

Bromine monoxide (BrO)

Bromine atoms (Br)

Chlorine atoms (Cl)

largest

reservoirs

most

 reactive

Halon-1301

HCFCs

Ultraviolet (UV)

sunlight &

other chemical

reactions

Conversion

H

Reactive halogen gases

alogen source gases

Ozone-depleting substances

Q8: What are the reactive halogen gases that destroy stratospheric ozone?

Emissions from human activities and natural processes are large sources of chlorine- and bromine-containing

gases for the stratosphere. When exposed to ultraviolet radiation from the Sun, these halogen source gases are

converted to more reactive gases also containing chlorine and bromine. Important examples of the reactive

gases that destroy stratospheric ozone are chlorine monoxide (ClO) and bromine monoxide (BrO). These and

other reactive gases participate in “catalytic” reaction cycles that efficiently destroy ozone. Volcanoes can emit

some chlorine-containing gases, but these gases are ones that readily dissolve in rainwater and ice and are

usually “washed out” of the atmosphere before they can reach the stratosphere.

Figure Q8-1. Conversion of halogen source gases. Halogen source gases (also known as ozone-depleting substances) are chemically converted to reactive halogen gases primarily in the stratosphere. The conversion requires ultraviolet sunlight and a few other chemical reactions. The short-lived gases undergo some conversion in the troposphere.

The reactive halogen gases contain all the chlorine and bromine originally present in the source gases. Some reactive

gases serve as reservoirs of chlorine and bromine whereas others participate in ozone destruction cycles.
TWENTY QUESTIONS

the reactive gases HCl, ClONO2, ClO, and other minor

gases. Available chlorine is constant within a few percent

from the surface to 47 kilometers (31 miles) altitude. In

the troposphere, available chlorine is contained almost

entirely in the source gases described in Figure Q7-1. At

higher altitudes, the source gases become a smaller fraction of available chlorine as they are converted to reactive

chlorine gases. At the highest altitudes, available chlorine

is all in the form of reactive chlorine gases.

In the altitude range of the ozone layer, as shown in

Figure Q8-2, the reactive chlorine gases HCl and

ClONO2 account for most of available chlorine. ClO, the

most reactive gas in ozone depletion, is a small fraction of

available chlorine. This small value limits the amount of

ozone destruction that occurs outside of polar regions.

Reactive chlorine in polar regions. Reactive chlorine gases in polar regions in summer look similar to the

altitude profiles shown in Figure Q8-2. In winter, however, the presence of polar stratospheric clouds (PSCs)

causes further chemical changes (see Q10). PSCs convert

HCl and ClONO2 to ClO when temperatures are near

minimum values in the winter Arctic and Antarctic stratosphere. In that case, ClO becomes the principal reactive

chlorine species in sunlit regions and ozone loss becomes

very rapid. An example of the late-winter ClO and ozone

distributions is shown in Figure Q8-3 for the Antarctic

stratosphere. These space-based measurements show that

ClO abundances are high in the lower stratosphere over a

region that exceeds the size of the Antarctic continent

(greater than 13 million square kilometers or 5 million

square miles). The peak abundance of ClO exceeds 1500

parts per trillion, which is much larger than typical midlatitude values shown in Figure Q8-2 and represents a

large fraction of reactive chlorine in that altitude region.

Because high ClO amounts cause rapid ozone loss (see

Q9), ozone depletion is found in regions of elevated ClO

(see Figure Q8-3).

Reactive bromine observations. Fewer measurements are available for reactive bromine gases than for

reactive chlorine, in part because of the low abundance of

bromine in the stratosphere. The most widely observed

bromine gas is bromine monoxide (BrO). The measured

abundances are consistent with values expected from the

conversion of the bromine-containing source gases such

as the halons and methyl bromide. 

Other sources. Some reactive halogen gases are

also produced at Earth’s surface by natural processes and

by human activities. However, because reactive halogen

gases are soluble in water, almost all become trapped in

the lower atmosphere by dissolving in rainwater and ice,

and ultimately are returned to Earth’s surface before they

can reach the stratosphere. For example, reactive chlorine is present in the atmosphere as sea salt (sodium chloride) produced by evaporation of ocean spray. Because

sea salt dissolves in water, this chlorine is removed and

does not reach the stratosphere in appreciable quantities.

Another ground-level source is emission of chlorine gases

from swimming pools, household bleach, and other uses.

When released to the atmosphere, this chlorine is rapidly

converted to forms that are soluble in water and removed.

The Space Shuttle and other rocket motors release reactive chlorine gases directly in the stratosphere. In that

case the quantities are very small in comparison with

other stratospheric sources. 

Measurements of Chlorine Gases from Space

Chlorine

monoxide (ClO)

Chlorine

nitrate (ClONO2)

Other gases

Altitude (kilometers)

0

10

20

30

40

4000

Hydrogen

chloride (HCl)

Chlorine source gases

(CFCs, HCFCs, carbon

tetrachloride, etc.) Available

chlorine

0 1000 2000 3000

Chlorine abundance (parts per trillion)Ozone Layer

Stratosphere Troposphere

Altitude (miles)

20

15

10

5

25

0

November 1994 (35°-49°N)

Figure Q8-2. Reactive chlorine gas observations.

Reactive chlorine gases and chlorine source gases can

be measured from space. These abundance measurements made over midlatitudes, for example, show that

chlorine-containing source gases are present in the troposphere and that reactive chlorine gases are present

in the stratosphere. In the stratosphere, reactive chlorine gases increase with altitude and chlorine source

gases decrease with altitude. This results because the

source gases are converted to reactive gases in chemical reactions involving ultraviolet sunlight. The principal

reactive gases formed are HCl, ClONO2, and ClO. The

sum of reactive gases and source gases gives available

chlorine, which is nearly constant with altitude up to 47

km. In the ozone layer, HCl and ClONO2 are the most

abundant reactive chlorine gases. (The unit “parts per

trillion” is defined in the caption of Figure Q7-1.)

Q.14
TWENTY QUESTIONS

Q.15

Volcanoes. Volcanic plumes generally contain large

quantities of chlorine in the form of hydrogen chloride

(HCl). Because the plumes also contain a considerable

amount of water vapor, the HCl is efficiently scavenged by

rainwater and ice and removed from the atmosphere. As a

result, most of the HCl in the plume does not enter the

stratosphere. After large recent eruptions, the increase in

HCl in the stratosphere has been small compared with the

total amount of chlorine in the stratosphere from other

sources.

Depleted ozone

Satellite Observations in the Lower Stratosphere

30 August 1996

Elevated chlorine

monoxide (ClO)

Global Ozone Dobson Network

The first instrument for routine monitoring of total ozone was developed by Gordon M. B. Dobson in

the 1920s. The instrument measures the intensity of sunlight at two ultraviolet wavelengths: one that is

strongly absorbed by ozone and one that is weakly absorbed. The difference in light intensity at the two

wavelengths is used to provide a measurement of total ozone above the instrument location. 

A global network of ground-based, total ozone observing stations was established in 1957 as part of

the International Geophysical Year. Today, there are about 100 sites distributed throughout the world (from

South Pole, Antarctica (90°S) to Ellesmere Island, Canada (83°N)), many of which routinely measure total

ozone with Dobson instruments. The accuracy of these observations is maintained by regular calibrations

and intercomparisons. Data from the network have been essential for understanding the effects of chlorofluorocarbons (CFCs) and other ozone-depleting gases on the global ozone layer, starting before the launch

of space-based ozone-measuring instruments and continuing to the present day. Because of their stability

and accuracy, the Dobson instruments are now routinely used to help calibrate space-based observations

of total ozone. 

Pioneering scientists have traditionally been honored by having units of measure named after them.

Accordingly, the unit of measure for total ozone is called the “Dobson unit” (see Q4). 

Figure Q8-3. Antarctic chlorine monoxide and

ozone. Satellite instruments monitor ozone and reactive

chlorine gases in the global stratosphere. Results are

shown here for Antarctic winter for a narrow altitude

region within the ozone layer. In winter, chlorine monoxide (ClO) reaches high values (1500 parts per trillion) in

the ozone layer, much higher than observed anywhere

else in the stratosphere because ClO is produced by

reactions on polar stratospheric clouds. These high ClO

values in the lower stratosphere last for 1 to 2 months,

cover an area that at times exceeds that of the Antarctic

continent, and efficiently destroy ozone in sunlit regions.

Ozone values measured simultaneously within the

ozone layer show very depleted values. 
TWENTY QUESTIONS

Q.16

Stratospheric ozone is destroyed by reactions involving reactive halogen gases, which are produced in the

chemical conversion of halogen source gases (see Figure

Q8-1). The most reactive of these gases are chlorine

monoxide (ClO) and bromine monoxide (BrO), and chlorine and bromine atoms (Cl and Br). These gases participate in three principal reaction cycles that destroy ozone. 

Cycle 1. Ozone destruction Cycle 1 is illustrated in

Figure Q9-1. The cycle is made up of two basic reactions:

ClO + O and Cl + O3. The net result of Cycle 1 is to convert one ozone molecule and one oxygen atom into two

oxygen molecules. In each cycle chlorine acts as a catalyst because ClO and Cl react and are reformed. In this

way, one Cl atom participates in many cycles, destroying

many ozone molecules. For typical stratospheric conditions at middle or low latitudes, a single chlorine atom can

destroy hundreds of ozone molecules before it reacts with

another gas to break the catalytic cycle.

Polar Cycles 2 and 3. The abundance of ClO is

greatly increased in polar regions during winter as a result

of reactions on the surfaces of polar stratospheric cloud

(PSC) particles (see Q10). Cycles 2 and 3 (see Figure

Q9-2) become the dominant reaction mechanisms for

polar ozone loss because of the high abundances of ClO

and the relatively low abundance of atomic oxygen

(which limits the rate of ozone loss by Cycle 1). Cycle 2

begins with the self-reaction of ClO. Cycle 3, which

begins with the reaction of ClO with BrO, has two reaction pathways to produce either Cl and Br or BrCl. The

net result of both cycles is to destroy two ozone molecules and create three oxygen molecules. Cycles 2 and 3

account for most of the ozone loss observed in the Arctic

and Antarctic stratospheres in the late winter/spring

season (see Q11 and Q12). At high ClO abundances, the

rate of ozone destruction can reach 2 to 3% per day in

late winter/spring.

Cl

Cl

Ozone Destruction Cycle 1

Chlorine atom (Cl) Ozone (O3)

Ozone

destruction

Chlorine monoxide (ClO)

Chlorine

catalytic

cycle

Oxygen molecule (O2)

Oxygen atom (O)

ClO + O

reaction

Cl + O3

reaction

ClO + O4 Cl + O2

Cl + O3 ClO + O

2

 O2

Net: O + O3 4 2O2

Oxygen molecule (O2)

3 4

Figure Q9-1. Ozone destruction

Cycle 1. The destruction of ozone in

Cycle 1 involves two separate chemical

reactions. The net or overall reaction is

that of atomic oxygen with ozone, forming two oxygen molecules. The cycle

can be considered to begin with either

ClO or Cl. When starting with ClO, the

first reaction is ClO with O to form Cl.

Cl then reacts with (and thereby

destroys) ozone and reforms ClO. The

cycle then begins again with another

reaction of ClO with O. Because Cl or

ClO is reformed each time an ozone

molecule is destroyed, chlorine is considered a catalyst for ozone destruction. Atomic oxygen (O) is formed

when ultraviolet sunlight reacts with

ozone and oxygen molecules. Cycle 1

is most important in the stratosphere at

tropical and middle latitudes where

ultraviolet sunlight is most intense.

Q9: What are the chlorine and bromine reactions that destroy stratospheric ozone? 

Reactive gases containing chlorine and bromine destroy stratospheric ozone in “catalytic” cycles made up of

two or more separate reactions. As a result, a single chlorine or bromine atom can destroy many hundreds of

ozone molecules before it reacts with another gas, breaking the cycle. In this way, a small amount of reactive

chlorine or bromine has a large impact on the ozone layer. Special ozone destruction reactions occur in polar

regions because the reactive gas chlorine monoxide reaches very high levels there in the winter/spring season.
TWENTY QUESTIONS

Q.17

Sunlight requirement. Sunlight is required to complete and maintain Cycles 1 through 3. Cycle 1 requires

sunlight because atomic oxygen is formed only with

ultraviolet sunlight. Cycle 1 is most important in the

stratosphere at tropical and middle latitudes where sunlight is most intense.

In Cycles 2 and 3, sunlight is required to complete the

reaction cycles and to maintain ClO abundances. In the

continuous darkness of winter in the polar stratospheres,

reaction Cycles 2 and 3 cannot occur. It is only in late

winter/spring when sunlight returns to the polar regions

that these cycles can occur. Therefore, the greatest destruction of ozone occurs in the partially to fully sunlit periods

after midwinters in the polar stratospheres. The sunlight

needed in Cycles 2 and 3 is not sufficient to form ozone

because ozone formation requires ultraviolet sunlight. In

the stratosphere in the winter/spring period, ultraviolet sunlight is weak because Sun angles are low. As a result,

ozone is destroyed in Cycles 2 and 3 in the sunlit winter

stratosphere but is not produced in significant amounts.

Other reactions. Atmospheric ozone abundances

are controlled by a wide variety of reactions that both produce and destroy ozone (see Q2). Chlorine and bromine

catalytic reactions are but one group of ozone destruction

reactions. Reactive hydrogen and reactive nitrogen gases,

for example, are involved in other catalytic ozonedestruction cycles that also occur in the stratosphere.

These reactions occur naturally in the stratosphere and

their importance has not been as strongly influenced by

human activities as have reactions involving halogens. 

ClO + ClO 4 (ClO)2

(ClO)2 + sunlight 4 ClOO + Cl

ClOO 4 Cl + O2

2(Cl + O3 4 ClO + O2)

Net: 2O3 4 3O2

ClO + BrO 4 Cl + Br + O2

BrCl + sunlight 4 Cl + Br 

Net: 2O3 4 3O2

ClO + BrO 4 BrCl + O2

Cl + O3 4 ClO + O2

Br + O3 4 BrO + O2

Cycle 2 Cycle 3

Ozone Destruction Cycles

) ) or

Figure Q9-2. Polar ozone destruction Cycles 2 and 3. Significant destruction of ozone occurs in polar regions

because ClO abundances reach large values. In that case, the cycles initiated by the reaction of ClO with another ClO

(Cycle 2) or the reaction of ClO with BrO (Cycle 3) efficiently destroy ozone. The net reaction in both cases is two ozone

molecules forming three oxygen molecules. The reaction of ClO with BrO has two pathways to form the Cl and Br product gases. Ozone destruction Cycles 2 and 3 are catalytic, as illustrated for Cycle 1 in Figure Q9-1, because chlorine

and bromine react and are reformed in each cycle. Sunlight is required to complete each cycle and to help form and

maintain ClO abundances.
TWENTY QUESTIONS

Q.18

The severe depletion of stratospheric ozone in

Antarctic winter is known as the “ozone hole” (see Q11).

Severe depletion first appeared over Antarctica because

atmospheric conditions there increase the effectiveness of

ozone destruction by reactive halogen gases (see Q8).

The formation of the Antarctic ozone hole requires abundant reactive halogen gases, temperatures low enough to

form polar stratospheric clouds (PSCs), isolation of air

from other stratospheric regions, and sunlight. 

Distributing halogen gases. Halogen source gases

emitted at Earth’s surface are present in comparable abundances throughout the stratosphere in both hemispheres

even though most of the emissions occur in the Northern

Hemisphere. The abundances are comparable because

most source gases have no important natural removal

processes in the lower atmosphere and because winds and

warm-air convection redistribute and mix air efficiently

throughout the troposphere. Halogen gases (in the form

of source gases and some reactive products) enter the

stratosphere primarily from the tropical troposphere.

Atmospheric air motions then transport them upward and

toward the poles in both hemispheres.

Low temperatures. The severe ozone destruction

represented by the ozone hole requires that low temperatures be present over a range of stratospheric altitudes,

over large geographical regions, and for extended time

periods. Low temperatures are important because

they allow polar stratospheric clouds (PSCs) to form.

Q10: Why has an “ozone hole” appeared over Antarctica when ozone-depleting gases are

present throughout the stratosphere? 

Ozone-depleting gases are present throughout the stratospheric ozone layer because they are transported great

distances by atmospheric air motions. The severe depletion of the Antarctic ozone layer known as the “ozone

hole” forms because of the special weather conditions that exist there and nowhere else on the globe. The very

cold temperatures of the Antarctic stratosphere create ice clouds called polar stratospheric clouds (PSCs).

Special reactions that occur on PSCs and the relative isolation of polar stratospheric air allow chlorine and

bromine reactions to produce the ozone hole in Antarctic springtime.

Nov Dec Jan Feb March April

40° to 90° Latitude

Arctic Winter

Minimum Air Temperatures in the 

Polar Lower Stratosphere

Arctic

Antarctic

-65

-70

-75

-80

-85

-90

-95

Temperature (degrees Celsius)

 PSC formation

 temperature

-140

-120

-100

-80

Temperature (degrees Fahrenheit)

-90

-110

-130

May June July August Sep Oct

Antarctic Winter

Range of Values 

Average winter values 

Arctic 1978-79 to 2001-02

Antarctic 1979 to 2001

-100

-150

-105

Figure Q10-1. Arctic and Antarctic temperatures. Stratospheric air temperatures

in both polar regions reach minimum values

in the lower stratosphere in the winter season. Minimum values over Antarctica reach

–90°C in July and August in an average

year. Over the Arctic, minimum average

values are near –80°C in January and

February. Polar stratospheric clouds

(PSCs) are formed when winter temperatures fall below the formation temperature

(about –78°C). This occurs on average for

weeks to months over both poles (see

heavy red and blue lines). Reactions on

PSCs cause the highly reactive chlorine gas

ClO to be formed, which increases the

destruction of ozone. The range of minimum temperatures found in the Arctic is

much greater than in the Antarctic. In some

years, PSC formation temperatures are not

reached in the Arctic, and significant ozone

depletion does not occur. In the Antarctic,

PSCs are formed over many months, and

severe ozone depletion now occurs in each

winter season.
TWENTY QUESTIONS

Reactions on the surfaces of the cloud particles initiate a

remarkable increase in the most reactive halogen gases

(see below and Q8). Temperatures are lowest in the stratosphere over both polar regions in winter. In the

Antarctic winter, minimum temperatures are generally

lower and less variable than in the Arctic winter (see

Figure Q10-1). Antarctic temperatures also remain

below the PSC formation temperature for much longer

periods during winter. This occurs, in part, because there

are significant differences between the hemispheres in the

distributions of land, ocean, and mountains at middle and

high latitudes. The winter temperatures are low enough

for PSCs to form for nearly the entire Antarctic winter but

usually only for part of every Arctic winter. 

Isolated conditions. Air in the polar stratospheric

regions is relatively isolated from other stratospheric

regions for long periods in the winter months. The isolation comes about because of strong winds that encircle

the poles, preventing substantial motion of air in or out of

the polar stratospheres. The isolation is much more effective in the Antarctic. Once chemical changes occur in the

low-temperature air as a result of the presence of PSCs,

the changes remain for many weeks to months.

Polar stratospheric clouds (PSCs). Polar stratospheric clouds cause changes in the abundances of reactive

chlorine gases. Reactions occur on the surfaces of PSC particles that convert the reservoir forms of reactive chlorine

gases, ClONO2 and HCl, to the most reactive form ClO (see

Figure Q8-1). ClO increases from a small fraction of available reactive chlorine gases to nearly all that is available

(see Q8). With increased ClO, additional catalytic cycles

involving ClO and BrO become active in the chemical

destruction of ozone when sunlight is available (see Q9). 

PSCs form when stratospheric temperatures fall

below about –78°C (–108°F) in polar regions. As a result,

PSCs are often found over vast areas of the winter polar

regions and over a significant altitude range. At low polar

temperatures, nitric acid (HNO3) and sulfur-containing

gases condense with water vapor to form solid and liquid

PSC particles. At even lower temperatures, ice particles

also form. PSC particles grow large enough and are

numerous enough that cloud-like features can be

observed from the ground under certain conditions, particularly when the Sun is near the horizon (see Figure

Q10-2). PSCs are often found near mountain ranges in

polar regions because the motion of air over the mountains can cause local cooling of stratospheric air.

When temperatures warm in spring, PSCs no longer

form and the production of ClO ends. Without continued

ClO production, ClO amounts decrease as other chemical

reactions reform ClONO2 and HCl. As a result, the

intense period of ozone depletion ends.

PSC removal. Once formed, PSC particles move

downward because of gravity. The largest particles move

down several kilometers or more in the stratosphere during the low-temperature winter/spring period. Because

most PSCs contain nitric acid, their downward motion

removes nitric acid from regions of the ozone layer. That

process is called denitrification. With less nitric acid, the

highly reactive chlorine gas ClO remains chemically

active for a longer period, thereby increasing chemical

ozone destruction. Because PSC formation temperatures

are required (see Figure Q10-1), denitrification occurs

each winter in the Antarctic and in some, but not all,

Arctic winters.

Q.19

Figure Q10-2. Polar stratospheric clouds. This photograph of an Arctic polar stratospheric cloud (PSC) was

taken from the ground at Kiruna, Sweden (67°N), on 27

January 2000. PSCs form during winters in the Arctic and

Antarctic stratospheres. The particles grow from the condensation of water along with nitrogen and sulfur gases.

Because the particles are large and numerous, the clouds

often can be seen with the human eye when the Sun is

near the horizon. Reactions on PSCs cause the highly

reactive chlorine gas ClO to be formed, which is very

effective in the chemical destruction of ozone.

Arctic Polar Stratospheric Clouds
TWENTY QUESTIONS

Q.20

Discovering the role of PSCs. The formation of

PSCs has been recognized for many years from groundbased observations. However, the geographical and altitude extent of PSCs in both polar regions was not known

fully until PSCs were observed by a satellite instrument in

the late 1970s. The role of PSCs in converting reactive

chlorine gases to ClO was not understood until after the

discovery of the Antarctic ozone hole in 1985. Our

understanding of the PSC role developed from laboratory

studies of their surface reactivity, computer modeling

studies of polar stratospheric chemistry, and direct sampling of PSC particles and reactive chlorine gases, such as

ClO, in the polar stratospheric regions.

The Discovery of the Antarctic Ozone Hole

The first decreases in Antarctic ozone were observed in the early 1980s over research stations

located on the Antarctic continent. The observations showed unusually low total overhead ozone during

the late winter/early spring months of September, October, and November. Total ozone was lower in these

months compared with previous observations made as early as 1957. The early-published reports came

from the British Antarctic Survey and the Japan Meteorological Agency. The results became more widely

known in the international community after they were published in the journal Nature in 1985 by three scientists from the British Antarctic Survey. Soon after, satellite measurements confirmed the spring ozone

depletion and further showed that in each late winter/spring season starting in the early 1980s, the depletion extended over a large region centered near the South Pole. The term “ozone hole” came about from

satellite images of total ozone that showed very low values encircling the Antarctic continent each spring

(see Q11). Currently, the formation and severity of the Antarctic “ozone hole” are documented each year by

a combination of satellite, ground-based, and balloon observations of ozone.
The severe depletion of Antarctic ozone known as the

“ozone hole” was first observed in the early 1980s. The

depletion is attributable to chemical destruction by reactive halogen gases, which increased in the stratosphere in

the latter half of the 20th century (see Q16). Conditions in

the Antarctic winter stratosphere are highly suitable for

ozone depletion because of (1) the long periods of

extremely low temperatures, which promote PSC formation and removal; (2) the abundance of reactive halogen

gases, which chemically destroy ozone; and (3) the isolation of stratospheric air during the winter, which allows

time for chemical destruction to occur (see Q10). The

severity of Antarctic ozone depletion can be seen using

images of total ozone from space, ozone altitude profiles,

and long-term average values of polar total ozone.

Antarctic ozone hole. The most widely used images

of Antarctic ozone depletion are those from space-based

measurements of total ozone. Satellite images made during

Antarctic winter and spring show a large region centered

near the South Pole in which total ozone is highly depleted

(see Figure Q11-1). This region has come to be called the

“ozone hole” because of the near-circular contours of low

ozone values in the images. The area of the ozone hole

has reached 25 million square kilometers (about 10 million square miles) in recent years, which is nearly twice

the area of the Antarctic continent. Minimum values of

total ozone inside the ozone hole have fallen as low as 100

Dobson units (DU) compared with normal springtime

values of about 300 DU (see Q4). The mass of ozone

destroyed over the Antarctic each season has reached an

estimated 80 megatons, which is about 3% of the global

ozone mass (1 megaton = 1 billion kilograms = 2.2 billion

pounds).

Altitude profiles of Antarctic ozone. Ozone within

the “ozone hole” is also measured using balloonborne

instruments (see Q5). Balloon measurements show changes

within the ozone layer, the vertical region that contains the

highest ozone abundances in the stratosphere. At geographic locations where the lowest total ozone values occur

in ozone hole images, balloon measurements show that the

chemical destruction of ozone is complete over a vertical

region of several kilometers. Figure Q11-2 shows an

example of such depletion with balloon measurements

TWENTY QUESTIONS

Q.21

III. STRATOSPHERIC OZONE DEPLETION

Q11: How severe is the depletion of the Antarctic ozone layer?

100 300 400 500

Total Ozone (Dobson units)

Antarctic Ozone Hole

4 October 2001

200

Figure Q11-1. Antarctic “ozone hole.” Total ozone

values are shown for high southern latitudes as measured by a satellite instrument. The dark regions over the

Antarctic continent show the severe ozone depletion now

found in every spring. Minimum values of total ozone

inside the ozone hole are close to 100 Dobson units (DU)

compared with normal springtime values of about 300

DU (see Q4). In late spring or early summer (NovemberDecember) the ozone hole disappears as ozone-depleted

air is displaced and diluted by ozone-rich air from outside the ozone hole.

Severe depletion of the Antarctic ozone layer was first observed in the early 1980s. Antarctic ozone depletion is

seasonal, occurring primarily in late winter and spring (August-November). Peak depletion occurs in October

when ozone is often completely destroyed over a range of altitudes, reducing overhead total ozone by as much

as two-thirds at some locations. This severe depletion creates the “ozone hole” in images of Antarctic total

ozone made from space. In most years the maximum area of the ozone hole usually exceeds the size of the

Antarctic continent.
made over South Pole, Antarctica, on 2 October 2001. The

altitude region of total depletion (14-20 kilometers) in the

profile corresponds to the region of lowest winter temperatures and highest ClO abundances. The average South Pole

ozone profiles for the decades 1962-1971 and 1992-2001

(see Figure Q11-2) show how reactive halogen gases have

changed the ozone layer. In the 1960s, the ozone layer is

clearly evident in the October average profile. In the 1990s,

minimum average values in the center of the layer have

fallen by 90% from the earlier values. 

Long-term total ozone changes. Low winter temperatures and isolated conditions occur each year in the

Antarctic stratosphere. As a result, significant spring ozone

depletion has been observed every year since the early

1980s. In prior years, the amounts of reactive halogen gases

in the stratosphere were insufficient to cause significant

depletion. Satellite observations can be used to examine

the average total ozone abundances in both polar regions

for the last three decades (see Figure Q12-1). In the

Antarctic, average values decreased steadily through the

1980s and 1990s, reaching minimum values that were 37%

less than in pre-ozone-hole years (1970-1982). The yearto-year changes in the average values reflect variations in

the meteorological conditions, which affect the extent of

low polar temperatures and the transport of air into and out

of the Antarctic winter stratosphere. However, essentially

all of the ozone depletion in the Antarctic in most years is

attributable to chemical loss from reactive halogen gases.

Restoring ozone in spring. The depletion of

Antarctic ozone occurs primarily in the late winter/spring

season. In spring, temperatures in the lower polar stratosphere eventually warm, thereby ending PSC formation

and the most effective chemical cycles that destroy ozone

(see Q10). The transport of air between the polar stratosphere and lower latitudes also increases during this time,

ending winter isolation. This allows ozone-rich air to be

transported to polar regions, displacing air in which ozone

has been severely depleted. This displaced air is diluted

at lower latitudes with more abundant ozone-rich air. As

a result, the ozone hole disappears by December and

Antarctic ozone amounts remain near normal until the

next winter season.

TWENTY QUESTIONS

Q.22

Figure Q11-2. Arctic and

Antarctic ozone distribution.

The stratospheric ozone layer

resides between about 10 and

50 kilometers (6 to 31 miles)

above Earth’s surface over the

globe. Long-term observations

of the ozone layer from small

balloons allow the winter

Antarctic and Arctic regions to

be compared. In the Antarctic

at the South Pole, halogen

gases have destroyed ozone in

the ozone layer beginning in

the 1980s. Before that period,

the ozone layer was clearly

present as shown here using

average ozone values from balloon observations made

between 1962 and 1971. In more recent years, as shown here for 2 October 2001, ozone is destroyed completely

between 14 and 20 kilometers (8 to 12 miles) in the Antarctic in spring. Average October values in the ozone layer now

are reduced by 90% from pre-1980 values. The Arctic ozone layer is still present in spring as shown by the average

March profile obtained over Finland between 1988 and 1997. However, March Arctic ozone values in some years are

often below normal average values as shown here for 30 March 1996. In such years, winter minimum temperatures are

generally below PSC formation temperatures for long periods. (Ozone abundances are shown here with the unit “milliPascals” (mPa), which is a measure of absolute pressure (100 million mPa = atmospheric sea-level pressure).)

0 5 10 15 0 5 10 15

0

5

10

15

20

25

30

Altitude (km)

Ozone abundance (mPa)

October averages

1962 - 1971

1992 - 2001

2 October 2001

March average

30 March 1996

1988 - 1997

Antarctic Ozone Arctic Ozone

Sodankyla,

Altitude (miles)

0

5

10

15

Polar Ozone Depletion

South Pole (90°S)

Finland (67°N)
Significant ozone depletion in the Arctic stratosphere

occurs in cold winters because of reactive halogen gases.

The depletion, however, is much less than occurs in every

Antarctic winter and spring. Although Arctic depletion

does not generally create persistent “ozone hole”-like features in Arctic total ozone maps, depletion is observed in

altitude profiles of ozone and in long-term average values

of polar ozone.

Altitude profiles of Arctic ozone. Arctic ozone is

measured using balloonborne instruments (see Q5), as in

the Antarctic (see Q11). Balloon measurements show

changes within the ozone layer, the vertical region that

contains the highest ozone abundances in the stratosphere.

Figure Q11-2 shows an example of a depleted ozone profile in the Arctic region on 30 March 1996, and contrasts

the depletion with that found in the Antarctic. The 30

March spring profile shows much less depletion than

the 2 October spring profile in the Antarctic. In general,

some reduction in the Arctic ozone layer occurs each

winter/spring season. However, complete depletion each

year over a broad vertical layer, as is now common in the

Antarctic stratosphere, is not found in the Arctic.

Long-term total ozone changes. Satellite observations can be used to examine the average total ozone abundances in the Arctic region for the last three decades and

to contrast them with results from the Antarctic (see

Figure Q12-1). Decreases from the pre-ozone-hole

average values (1970-1982) were observed in the Arctic

beginning in the 1980s, when similar changes were

occurring in the Antarctic. The decreases have reached a

maximum of 22%, but have remained smaller than those

found in the Antarctic since the mid-1980s. The year-toyear changes in the average Arctic and Antarctic average

ozone values reflect annual variations in meteorological

conditions that affect the extent of low polar temperatures

and the transport of air into and out of the polar stratosphere. The effect of these variations is generally greater

for the Arctic than the Antarctic. In almost all years, most

of the Arctic ozone decrease (about 75%) is attributable

to chemical destruction by reactive halogen gases.

Arctic vs. Antarctic. The Arctic winter stratosphere

is generally warmer than its Antarctic counterpart (see

Figure Q10-1). Higher temperatures reduce polar stratospheric cloud (PSC) formation, the conversion of reacTWENTY QUESTIONS

Q.23

Q12: Is there depletion of the Arctic ozone layer?

Yes, significant depletion of the Arctic ozone layer now occurs in some years in the late winter/spring period

(January-April). However, the maximum depletion is generally less severe than that observed in the Antarctic

and is more variable from year to year. A large and recurrent “ozone hole,” as found in the Antarctic stratosphere, does not occur in the Arctic.

Figure Q12-1. Average polar ozone. Total

ozone in polar regions is measured by wellcalibrated satellite instruments. Shown here

is a comparison of average springtime totalozone values found between 1970 and 1982

(solid red line) with those in later years. Each

point represents a monthly average in

October in the Antarctic or in March in the

Arctic. After 1982, significant ozone depletion is found in most years in the Arctic and

all years in the Antarctic. The largest average

depletions have occurred in the Antarctic in

the last decade. Natural variations in meteorological conditions influence the year-toyear changes in depletion, particularly in the

Arctic. Essentially all of the decrease in the

Antarctic and usually most of the decrease

in the Arctic each year are attributable to

chemical destruction by reactive halogen

gases. Average total ozone values over the Arctic are initially larger each winter/spring season because more ozone is

transported poleward in the Northern Hemisphere each season than in the Southern Hemisphere.

500

450

400

350

300

250

200

1970 1975 1980 1985 1990 1995 2000

Year

Average Total Ozone in Polar Regions

Total ozone (Dobson units)

Chemical destruction

and natural variation

of ozone

1970-1982 average

(63°-90° average)

Arctic Region in March

Antarctic Region in October
tive chlorine gases to form ClO, and, as a consequence,

the amount of ozone depletion (see Q10). Furthermore,

the temperature and wind conditions are much more variable in the Arctic from winter to winter and within a winter

season than in the Antarctic. Large year-to-year differences occur in Arctic minimum temperatures and the duration of PSC-forming temperatures into early spring. In a

few Arctic winters, minimum temperatures are not low

enough for PSCs to form. These factors combine to cause

ozone depletion to be variable in the Arctic from year to

year, with some years having little to no ozone depletion. 

As in the Antarctic, depletion of ozone in the Arctic

is confined to the late winter/spring season. In spring,

temperatures in the lower stratosphere eventually warm,

thereby ending PSC formation and the most effective

chemical cycles that destroy ozone. The subsequent transport of ozone-rich air into the Arctic stratosphere displaces

ozone-depleted air. As a result, ozone layer abundances

are restored to near-normal values until the following

winter.

High Arctic total ozone. A significant difference

exists between the Northern and Southern Hemispheres

in how ozone-rich stratospheric air is transported into the

polar regions from lower latitudes during fall and winter.

In the northern stratosphere, the poleward and downward

transport of ozone-rich air is stronger. As a result, before

the onset of ozone depletion in the 1980s, total ozone

values in the Arctic were considerably higher than in the

Antarctic in winter and spring (see Figure Q12-1). The

Arctic values have remained higher to the present day

because of the greater depletion of Antarctic ozone.

TWENTY QUESTIONS

Q.24

Replacing the Loss of “Good” Ozone in the Stratosphere

The idea is sometimes put forth that humans could replace the loss of global stratospheric ozone,

called “good” ozone, by making ozone and transporting it to the stratosphere. Ozone amounts reflect a balance in the stratosphere between continual production and destruction by mostly naturally occurring reactions (see Q2). The addition of chlorine and bromine to the stratosphere from human activities has increased

ozone destruction and lowered “good” ozone amounts. Adding manufactured ozone to the stratosphere

would upset the existing balance. As a consequence, most added ozone would be destroyed in chemical

reactions within weeks to months as the balance was restored. So, it is not practical to consider replacing

the loss of global stratospheric ozone because the replacement effort would need to continue indefinitely, or

as long as increased chlorine and bromine amounts remained.

Other practical difficulties in replacing stratospheric ozone are the large amounts of ozone required

and the delivery method. The total amount of atmospheric ozone is approximately 3000 megatons (1

megaton = 1 billion kilograms) with most residing in the stratosphere. The replacement of the average

global ozone loss of 3% would require 90 megatons of stratospheric ozone to be distributed throughout the

layer located many kilometers above Earth’s surface. The energy required to produce this amount of ozone

would be a significant fraction of the electrical power generated in the United States, which is now approximately 5 trillion kilowatt hours. Processing and storing requirements for ozone, which is explosive and toxic

in large quantities, would increase the energy requirement. In addition, methods suitable to deliver and distribute large amounts of ozone to the stratosphere have not been demonstrated yet. Concerns for a global

delivery system would include further significant energy use and unforeseen environmental consequences.
Stratospheric ozone has decreased over the globe since

the 1980s. The depletion, which in the period 1997-2001

averaged about 3% (see Figure Q13-1), is larger than

natural variations in ozone. The observations shown in

Figure Q13-1 have been smoothed to remove regular

ozone changes that are due to seasonal and solar effects

(see Q14). The increase in reactive halogen gases in the

stratosphere is considered to be the primary cause of the

average depletion. The lowest ozone values in recent years

occurred following the 1991 eruption of Mt. Pinatubo,

which increased the number of sulfur-containing particles

in the stratosphere. The particles remain in the stratosphere

for several years, increasing the effectiveness of reactive

halogen gases in destroying ozone (see Q14).

Observed ozone depletion varies significantly with latitude on the globe (see Figure Q13-1). The largest losses

occur at the highest southern latitudes as a result of the severe

ozone loss over Antarctica each winter/spring period. The

next largest losses are observed in the Northern Hemisphere,

caused in part by winter/spring losses over the Arctic. Air

depleted in ozone over both polar regions spreads away from

the poles during and after each winter/spring period. Ozone

depletion also occurs directly at latitudes between the

equator and polar regions, but is smaller because of smaller

amounts of reactive halogen gases present there.

Tropical regions. There has been little or no depletion

of total ozone in the tropics (between about 20° latitude north

and south of the equator in Figure Q13-1). In this region of

the lower stratosphere, air has only recently (less than 18

months) been transported from the lower atmosphere. As a

result, the conversion of halogen source gases to reactive

halogen gases is very small. Because of the low abundance

of reactive gases, total ozone depletion in this region is also

very small. In contrast, stratospheric air in polar regions

has been in the stratosphere for an average of 4 to 7 years,

and the abundance of reactive halogen gases is much larger. 

Seasonal changes. The magnitude of global ozone

depletion also depends on season of the year. In the period

pre-1980 to 1997-2001, average total ozone decreased by

about 3% in northern middle latitudes (35°N-60°N) and

about 6% at southern middle latitudes (35°S-60°S). The

seasonality of these changes is different in the two hemispheres. In the Northern Hemisphere, larger decreases

are observed in winter/spring (4%) than in summer/-

autumn (2%). In the Southern Hemisphere, the decreases

are about the same (6%) during all seasons.

TWENTY QUESTIONS

Latitude

-90 -60 -30 0 30 60 90

0

-10

Ozone change (%)

-20

South North

1965 1970 1975 1980 1985 1990 1995 2000

Year

-4

-2

0

Average change

Uncertainty range

Ozone change (%)

Changes from 1964-1980 average

Changes between 1980 and 2000

Average change

Uncertainty range

Global Total Ozone Change

Figure Q13-1. Global total ozone changes. Global

total ozone values decreased by an average of a few

percent in the last two decades, as measured by satellite instruments. In the top panel, global ozone changes

are compared with average global ozone found in the

period of 1964 to 1980. Between 1980 and 2000, the

largest decreases occurred following the volcanic eruption of Mt. Pinatubo in 1991. In the 1997 to 2001 period

global ozone was reduced by about 3% from the 1964-

1980 average. In the bottom panel, ozone changes

between 1980 and 2000 are compared for different latitudes. The largest decreases have occurred at the

highest latitudes in both hemispheres because of the

large winter/spring depletion in polar regions. The

losses in the Southern Hemisphere are greater than

those in the Northern Hemisphere because of the

greater losses that occur each year in the Antarctic stratosphere. Long-term changes in the tropics are much

smaller because reactive halogen gases are not abundant in the tropical lower stratosphere.

Q13: How large is the depletion of the global ozone layer?

The ozone layer has been depleted gradually since 1980 and now is about an average of 3% lower over the

globe. The depletion, which exceeds the natural variations of the ozone layer, is very small near the equator

and increases with latitude toward the poles. The large average depletion in polar regions is primarily a result

of the late winter/spring ozone destruction that occurs there annually.

Q.25
Changes in solar radiation and increases in stratospheric particles from volcanic eruptions both affect the

abundance of stratospheric ozone, but they have not

caused the long-term decreases observed in total ozone.

Solar changes. The formation of stratospheric ozone

is initiated by ultraviolet (UV) radiation coming from the

Sun (see Figure Q2-1). As a result, an increase in the

Sun’s radiation output increases the amount of ozone in

Earth’s atmosphere. The Sun’s radiation output and

sunspot number vary over the well-known 11-year solar

cycle. Observations over several solar cycles (since the

1960s) show that global total ozone levels vary by 1 to

2% between the maximum and minimum of a typical

cycle. Changes in solar output at a wavelength of 10.7

cm, although much larger than changes in total solar

output, are often used to show when periods of maximum

and minimum total output occur (see Figure Q14-1).

Since 1978, the Sun’s output has gone through maximum

values around 1969, 1980, and 1991, and is currently near

a maximum value in 2002.

Over the last two decades, average total ozone has

decreased over the globe. Average values in recent years

TWENTY QUESTIONS

Q.26

Global Ozone, Volcanic Eruptions,

and the Solar Cycle 300

250

200

150

100

50

Incoming solar radiation at

10.7-cm wavelength

2

0

-2

-4

-6 (uncorrected for solar

 and seasonal effects)

Change from 1964-1980

average (%)

10.7-cm solar radiation

0.95

0.90

0.85

Tr

0.80

ansmission

1970 1980 1990 2000

Year

Solar transmission

Hawaii (20°N)

Volcanic eruptions

Total ozone 90°S-90°N

Figure Q14-1. Solar changes and volcanoes. Total

ozone values have decreased beginning in the early

1980s (see middle panel). The ozone values shown have

not been smoothed for solar and seasonal effects as they

were in Figure Q13-1. Incoming solar radiation, which

produces ozone in the stratosphere, changes on a wellrecognized 11-year cycle. Solar radiation at 10.7-cm

wavelength is often used to show the times of maximum

and minimum solar output (see top panel). A comparison of the top and middle panels indicates that the cyclic

changes in solar output cannot account for the long-term

decreases in ozone. Volcanic eruptions occurred frequently in the 1965 to 2002 period. The largest recent

eruptions are El Chichón (1982) and Mt. Pinatubo (1991)

(see large red arrows). Large volcanic eruptions are

monitored by the decreases in solar transmission to

Earth’s surface that occur because new particles are

formed in the stratosphere from volcanic sulfur emissions

(see bottom panel). These particles increase ozone

depletion but do not remain in the stratosphere for more

than a few years. A comparison of the middle and bottom

panels indicates that large volcanic eruptions also cannot

account for the long-term decreases found in global total

ozone.

Q14: Do changes in the Sun and volcanic eruptions affect the ozone layer?

Yes, factors such as changes in solar radiation, as well as the formation of stratospheric particles after volcanic

eruptions, do influence the ozone layer. However, neither factor can explain the average decreases observed in

global total ozone over the last two decades. If large volcanic eruptions occur in the coming decades, ozone

depletion will increase for several years after the eruption.
show a 3-4% depletion from pre-1980 values (see Figure

Q14-1). The ozone values shown have not been smoothed

for solar and seasonal effects as they were for Figure Q13-1.

Over the same period, changes in solar output show the

expected 11-year cycle but do not show a decrease with

time. For this reason, the long-term decreases in global

ozone cannot result from changes in solar output alone.

Most discussions of long-term ozone changes presented

in this and previous international scientific assessments

account for the influence of the 11-year solar cycle.

Past volcanoes. Large volcanic eruptions inject sulfur

gases directly into the ozone layer, causing new sulfate

particles to be formed. The particles initially form in the

stratosphere above the volcano location and then spread

globally as air is transported within the stratosphere. The

presence of volcanic particles in the stratosphere is shown

by observations of solar transmission through the atmosphere. When large amounts of particles are present in the

stratosphere, transmission of solar radiation is reduced.

The large eruptions of El Chichón (1982) and Mt. Pinatubo

(1991) are recent examples of events that temporarily

reduced transmission (see Figure Q14-1). 

Laboratory measurements and stratospheric observations have shown that chemical reactions on the surface

of volcanically produced particles increase ozone destruction by increasing the amounts of the highly reactive chlorine gas, chlorine monoxide (ClO). Ozone depletion

increases as a consequence of increased ClO. The most

recent large eruption was that of Mt. Pinatubo, which

resulted in up to a 10-fold increase in the number of particles available for surface reactions. Both El Chichón and

Mt. Pinatubo reduced global ozone for a few years (see

Figure Q14-1). After a few years, the effect of volcanic

particles diminishes as volcanically produced particles

are gradually removed from the stratosphere by natural

air circulation. Because of particle removal, the few large

volcanic eruptions of the last two decades cannot account

for the long-term decreases observed in ozone over the

same period. 

Future volcanoes. Observations and atmospheric

models indicate that the record-low ozone levels observed

in 1992-1993 resulted from the relatively large number of

particles produced by the Mt. Pinatubo eruption, combined with the relatively large amounts of halogen gases

present in the stratosphere in the 1990s. If the Mt.

Pinatubo eruption had occurred before 1980, changes to

global ozone would have been much smaller than

observed in 1992-1993 because the abundance of halogen

gases in the stratosphere was smaller. In the early decades

of the 21st century, the abundance of halogen gases will

still be substantial in the global atmosphere (see Figure

Q16-1). If large volcanic eruptions occur in these early

decades, ozone depletion will increase for several years.

If an eruption larger than Mt. Pinatubo occurs then ozone

losses could be larger than previously observed and persist longer. Only later in the 21st century when halogen

gas abundances have declined will the effect of volcanic

eruptions on ozone be lessened. 

TWENTY QUESTIONS

Q.27
Montreal Protocol. In 1985, a treaty called the

Convention for the Protection of the Ozone Layer was

signed by 20 nations in Vienna. The signing nations

agreed to take appropriate measures to protect the ozone

layer from human activities. The Vienna Convention

supported research, exchange of information, and future

protocols. In response to growing concern, the Montreal

Protocol on Substances that Deplete the Ozone Layer was

signed in 1987 and ratified in 1989. The Protocol established legally binding controls for developed and developing nations on the production and consumption of

halogen source gases known to cause ozone depletion.

National consumption of a halogen gas is defined as the

amount that production and imports of a gas exceed its

export to other nations. 

Amendments and Adjustments. As the scientific

basis of ozone depletion became more certain after 1987

and substitutes and alternatives became available for the

principal halogen source gases, the Montreal Protocol was

strengthened with Amendments and Adjustments. These

revisions added new controlled substances, accelerated

existing control measures, and scheduled phaseouts of the

production of certain gases. The initial Protocol called for

only a slowing of chlorofluorocarbon (CFC) and halon production. The 1990 London Amendments to the Protocol

called for a phaseout of the production of the most damaging ozone-depleting substances in developed nations

by 2000 and in developing nations by 2010. The 1992

Copenhagen Amendments accelerated the date of the

phaseout to 1996 in developed nations. Further controls on

ozone-depleting substances were agreed upon in later meetings in Vienna (1995), Montreal (1997), and Beijing (1999). 

Montreal Protocol projections. The future stratospheric abundances of effective stratospheric chlorine can

be projected based on the provisions of the Montreal

Protocol. The concept of effective stratospheric chlorine

accounts for the combined effect on ozone of chlorine- and

bromine-containing gases. The results are shown in Figure

Q15-1 for the following cases: 

• No Protocol and continued production increases of

3% per year (business-as-usual scenario). 

• Continued production as allowed by the Protocol’s

original provisions agreed upon in Montreal in

1987.

• Restricted production as outlined in the subsequent

Amendments and Adjustments: London, 1990;

Copenhagen, 1992; and Beijing, 1999. (The city

names and years signify where and when the agreements were made.)

• Zero emissions of ozone-depleting gases starting

in 2003. 

In each case, production of a gas is assumed to result

in its eventual emission to the atmosphere. Without the

Montreal Protocol, continued production and use of CFCs

and other ozone-depleting gases are projected to have

increased effective stratospheric chlorine tenfold by the

mid-2050s compared with the 1980 value. Such high

values likely would have increased global ozone depletion far beyond that currently observed. As a result,

harmful UV-B radiation would have increased substantially at Earth’s surface, causing a rise in excess skin

cancer cases (see lower panel of Figure Q15-1 and Q17). 

The 1987 provisions of the Montreal Protocol would

have only slowed the approach to high effective chlorine

values by one or more decades in the 21st century. Not

until the 1992 Copenhagen Amendments and Adjustments

did the Protocol projections show a decrease in future

effective stratospheric chlorine values. Now, with full

compliance to the Montreal Protocol and its Amendments

and Adjustments, use of the major human-produced

ozone-depleting gases will ultimately be phased out and

effective stratospheric chlorine will slowly decay,

reaching pre-ozone-hole values in the mid-21st century.

Zero emissions. Effective chlorine values in the

coming decades will be influenced by emissions of

halogen source gases produced in those decades as well

as the emission of currently existing gases that are now

being used or stored in various ways. Some continued

TWENTY QUESTIONS

Q.28

IV. CONTROLLING OZONE-DEPLETING GASES

Q15: Are there regulations on the production of ozone-depleting gases?

Yes, the production of ozone-depleting gases is regulated under a 1987 international agreement known as the

“Montreal Protocol on Substances that Deplete the Ozone Layer” and its subsequent Amendments and

Adjustments. The Protocol, now ratified by over 180 nations, establishes legally binding controls on the

national production and consumption of ozone-depleting gases. Production and consumption of all principal

halogen-containing gases by developed and developing nations will be significantly reduced or phased out

before the middle of the 21st century. 
production and consumption of ozone-depleting gases is

allowed, particularly in developing nations, under the

1999 Beijing agreements. As a measure of the contribution of these continued emissions to the effective chlorine

value, the “zero emissions” case is included in Figure

Q15-1. In this hypothetical case, all emissions of ozonedepleting gases are set to zero beginning in 2003. The

reductions in effective stratospheric chlorine beyond the

values expected for the 1999 Beijing agreement would be

significantly smaller than most earlier changes, as shown. 

HCFC substitute gases. The Montreal Protocol provides for the transitional use of hydrochlorofluorocarbons

(HCFCs) as substitute compounds for principal halogen

source gases such as CFC-12. HCFCs differ chemically

from most other halogen source gases in that they contain

hydrogen (H) atoms in addition to chlorine and fluorine

atoms. HCFCs are used for refrigeration, for blowing

foams, and as solvents, which were primary uses of CFCs.

HCFCs are 1 to 15% as effective as CFC-12 in depleting

stratospheric ozone because they are chemically removed

primarily in the troposphere. This removal partially protects stratospheric ozone from the halogens contained in

HCFCs. In contrast, CFCs and many other halogen source

gases are chemically inert in the troposphere and, hence,

reach the stratosphere without being significantly

removed. Because HCFCs still contribute to the halogen

abundance in the stratosphere, the Montreal Protocol

requires the production and consumption of HCFCs to

end in developed and developing nations by 2040.

HFC substitute gases. Hydrofluorocarbons (HFCs)

are also used as substitute compounds for CFCs and other

halogen source gases. HFCs contain only hydrogen,

fluorine, and carbon atoms. Because HFCs contain no

chlorine or bromine, they do not contribute to ozone

depletion. As a consequence, the HFCs are not regulated

by the Montreal Protocol. However, HFCs (as well as all

halogen source gases) are radiatively active gases that

contribute to human-induced global warming and climate

change as they accumulate in the atmosphere (see Q18).

HFCs are included in the group of gases listed in the Kyoto

Protocol of the United Nations Framework Convention

on Climate Change (UNFCC).

TWENTY QUESTIONS

0

10000

No Protocol

Predicted abundance (parts per trillion) 5000

15000

20000

Montreal

1987

London

1990

Beijing

1999

Copenhagen

 1992

Year

1980 2000 2020 2040 2060 2080 2100

0

100

200

300

400

500

Cases per million people per year

Excess skin

cancer cases

No Protocol

Montreal

1987

London

1990

Copenhagen

1992

Effect of the Montreal Protocol

Zero Emissions

Effective stratospheric

chlorine

Figure Q15-1. Effect of the Montreal Protocol. The

purpose of the Montreal Protocol is to achieve reductions

in stratospheric abundances of chlorine and bromine.

The reductions follow from restrictions on the production

and consumption of manufactured halogen source gases.

Predictions for the future abundance of effective stratospheric chlorine are shown in the top panel assuming (1)

no Protocol regulations, (2) only the regulations in the

original 1987 Montreal Protocol, and (3) additional regulations from the subsequent Amendments and

Adjustments. The city names and years indicate where

and when changes to the original 1987 Protocol provisions were agreed upon. Effective stratospheric chlorine

as used here accounts for the combined effect of chlorine and bromine gases. Without a Protocol, stratospheric halogen gases are projected to have increased

significantly in the 21st century. The “zero emissions” line

shows stratospheric abundances if all emissions were

reduced to zero beginning in 2003. The lower panel

shows how excess skin cancer cases (see Q17) would

increase with no regulations and how they will be reduced

under the Protocol provisions. (The unit “parts per trillion” is defined in the caption of Figure Q7-1.)

Q.29
Effective stratospheric chlorine. The Montreal

Protocol has been successful in slowing and reversing the

increase of ozone-depleting gases in the atmosphere. An

important measure of its success is the change in the value

of effective stratospheric chlorine. Effective chlorine is

based on measured or estimated abundances of both chlorine- and bromine-containing gases in the stratosphere. The

two groups of gases are not simply added together, because

bromine gases are much more effective on a per-atom basis

than chlorine in depleting ozone (see Q7). Instead, the

amounts for bromine gases are multiplied by a factor that

accounts for their greater effectiveness and then added to

the total amount of chlorine. Bromine atoms are much less

abundant in the stratosphere than chlorine (see Figure Q7-

1), but are about 45 times more effective than chlorine in

chemically destroying ozone molecules. Increases in effective chlorine in the past decades have caused stratospheric

ozone depletion. Accordingly, ozone is expected to recover

in the future as effective chlorine values decrease. 

Effective stratospheric chlorine changes. In the latter

half of the 20th century up until the 1990s, effective chlorine

values steadily increased (see Figure Q16-1). The values

shown were calculated using halogen source gas abundances

obtained from measurements, historical estimates of abundance, and future projections of abundance. As a result of

the Montreal Protocol regulations, the long-term increase in

effective chlorine slowed, reached a peak, and began to

decrease in the 1990s. This small and continuing decrease

means that the potential for stratospheric ozone depletion

has begun to lessen as a result of the Montreal Protocol. The

decrease in effective chlorine is projected to continue

throughout the 21st century if all nations continue to comply

with the provisions of the Protocol. Once halogen gas emissions from human activities have ceased, the decrease in

effective chlorine will continue as natural destruction

processes gradually reduce halogen gas abundances in the

global atmosphere. Significant reduction requires decades

because the lifetimes of halogen source gas molecules in

the atmosphere range up to 100 years (see Table Q7-1). 

Individual halogen source gases. Reductions in effective chlorine follow directly from decreases in the emission

of individual halogen source gases. The regulation of humanproduced halogen source gases under the Montreal Protocol

is considered separately for each class of one or more gases

and is based on several factors. The factors include (1) the

effectiveness of each class in depleting ozone in comparison to other halogen source gases; (2) the availability of

suitable substitute gases for domestic and industrial use; and

(3) the impact of regulation on developing nations.

The change in atmospheric abundance of a gas in

response to regulation also depends on a number of factors

that include (1) how the gas is used and released to the atmosphere; (2) how rapidly the gas is chemically destroyed in

the atmosphere; and (3) the total amount of the gas that has

accumulated in the atmosphere.

Methyl chloroform and CFCs. The largest reduction

in the abundance of a halogen source gas has occurred for

methyl chloroform (CH3CCl3). As shown in Figure Q16-1,

atmospheric methyl chloroform values dropped abruptly as a

result of the provisions of the Montreal Protocol that reduced

global production to near zero. Atmospheric abundances subsequently dropped rapidly because methyl chloroform has a

short atmospheric lifetime (about 5 years). In addition, methyl

chloroform is used mainly as a solvent and therefore has no

significant long-term storage reservoir, as do refrigerants, for

example. The reduction in effective chlorine in the 1990s

came primarily from the reduction in methyl chloroform

abundance in the atmosphere. Significant emissions reductions have also occurred for the chlorofluorocarbons CFC-11,

CFC-12, and CFC-113 starting in the 1990s. As a result, the

atmospheric increases of these gases have slowed, and CFC11 and CFC-113 abundances have decreased slightly (see

Figure Q16-1). Because of longer lifetimes in the atmosphere (see Table Q7-1), CFC abundances decrease more

slowly than methyl chloroform as their emissions are reduced. 

HCFC substitute gases. The Montreal Protocol allows

for the use of hydrochlorofluorocarbons (HCFCs) as shortterm substitutes for CFCs. As a result, the abundances of

HCFC-22, HCFC-141b, and HCFC-142b continue to grow

in the atmosphere (see Figure Q16-1). HCFCs are desirable

as CFC substitutes because they are partially destroyed in the

troposphere by chemical processes, thus reducing the overall

effectiveness of their emissions in destroying stratospheric

ozone. Under the Montreal Protocol, HCFC production will

reach zero in developed nations by 2030 and in developing

nations by 2040. Thus, the future projections in Figure Q16-

1 show HCFC abundances reaching a peak in the first decades

of the 21st century and steadily decreasing thereafter.

TWENTY QUESTIONS

Q.30

Q16: Has the Montreal Protocol been successful in reducing ozone-depleting gases in the

atmosphere?

Yes, as a result of the Montreal Protocol, the total abundance of ozone-depleting gases in the atmosphere has

begun to decrease in recent years. If the nations of the world continue to follow the provisions of the Montreal

Protocol, the decrease will continue throughout the 21st century. Some individual gases such as halons and

hydrochlorofluorocarbons (HCFCs) are still increasing in the atmosphere, but will begin to decrease in the next

decades if compliance with the Protocol continues. By midcentury, the effective abundance of ozone-depleting

gases should fall to values present before the Antarctic “ozone hole” began to form in the early 1980s.
Halons. The atmospheric abundances of Halon-1211

and Halon-1301 are significant fractions of all brominecontaining source gases (see Figure Q7-1) and continue to

grow despite the elimination of production in developed

nations in 1994 (see Figure Q16-1). The growth in abundance continues because substantial reserves are held in fireextinguishing equipment and are gradually being released,

and production and consumption are still allowed in developing nations. Release of stored halons could keep atmospheric halon abundances high well into the 21st century.

Methyl chloride and methyl bromide. Both methyl

chloride (CH3Cl) and methyl bromide (CH3Br) are unique

among halogen source gases because a substantial fraction

of their sources are associated with natural processes (see

Q7). The average atmospheric abundance of methyl chloride, which is not regulated under the Montreal Protocol,

is expected to remain nearly constant throughout this century. At century’s end, methyl chloride is expected to

account for a large fraction of remaining effective stratospheric chlorine because the abundances of other gases,

such as the CFCs, are expected to be greatly reduced (see

Figure Q16-1). The abundance of methyl bromine, which

is regulated under the Protocol, is projected to decrease in

the first decades of this century as a result of production

phaseouts in developed and developing countries. In the

remaining decades of the century, methyl bromide abundances are expected to be nearly constant.

TWENTY QUESTIONS

Q.31

Estimates of

historical abundance

Effective stratospheric

chlorine (parts per trillion)

Observations

Future projections

Atmospheric abundance (parts per trillion)

400

200

CFC-12

CFC-11

CFC-113

0

160

120

80

40

HCFC-22

CH3CCl3

CCl4

HCFC-141b

Halon-1211 Halon-1301

4

3

2

1

0

0

HCFC-142b

Year 1950 2000 2050 2100

0

3000

1980 Level

2500

2000

1500

1000

500

Effective Chlorine

500

400

8

10

CH3Cl

CH3Br

Past and Future Abundance of 

Atmospheric Halogen Source Gases

Figure Q16-1. Halogen source gas changes. The rise in

effective stratospheric chlorine values in the 20th century has

slowed and reversed in the last decade (top panel). Effective

chlorine values combine the measured or projected abundances of chlorine-containing gases with those of brominecontaining gases in a way that properly accounts for the

greater effectiveness of bromine in depleting stratospheric

ozone. As effective chlorine decreases in the 21st century,

the potential for ozone depletion from halogen gases will also

decrease. The decrease in effective chlorine values is a result

of reductions in individual halogen source gas emissions.

The emissions decreased because of the Montreal Protocol,

which restricts production and consumption of manufactured

halogen gases. The changes in the atmospheric abundance

of individual gases are shown in the lower panels using a

combination of direct atmospheric measurements, estimates

of historical abundance, and future projections of abundance.

The increases of CFCs, along with those of CCl4 and

CH3CCl3, have either slowed significantly or reversed in the

last decade. HCFCs, which are being used as CFC substitutes, will continue to increase in the coming decades. Some

halon abundances will also grow in the future while current

halon reserves are being depleted. Smaller relative

decreases are expected for CH3Br in response to restrictions

because it has substantial natural sources. CH3Cl has large

natural sources and is not regulated under the Montreal

Protocol. (See Figure Q7-1 for chemical names and formulas. The unit “parts per trillion” is defined in the caption of

Figure Q7-1.) 
The depletion of stratospheric ozone leads to an

increase in surface ultraviolet radiation. The increase

occurs primarily in the ultraviolet-B (UV-B) component

of the Sun’s radiation. UV-B is defined as radiation in the

wavelength range of 280 to 315 nanometers. Changes in

UV-B at the surface have been observed directly and can

be estimated from ozone changes.

Surface UV-B radiation. The amount of ultraviolet

radiation reaching Earth’s surface depends in large part

on the amount of ozone in the atmosphere. Ozone molecules in the stratosphere absorb UV-B radiation, thereby

significantly reducing the amount of this radiation that

reaches Earth’s surface (see Q3). If total ozone amounts

are reduced in the stratosphere, then the amount of UV

radiation reaching Earth’s surface generally increases.

This relationship between total ozone and surface UV

radiation has been studied at a variety of locations with

direct measurements of both ozone and UV. The actual

amount of UV reaching a location depends on a large

number of additional factors, including the position of the

Sun in the sky, cloudiness, and air pollution. In general,

surface UV at a particular location on Earth changes

throughout the day and with season as the Sun’s position

in the sky changes.

Long-term surface UV changes. Satellite observations of long-term global ozone changes can be used to

estimate changes in global surface UV that have occurred

over the past two decades. These changes are of interest

because UV radiation can cause harm to humans and other

life forms (see Q3). The amount of UV that produces an

“erythemal” or sunburning response in humans is often

separately evaluated. UV-B radiation is a large component of sunburning UV. Long-term changes in sunburning

UV at a location can be estimated from the changes in

total ozone at that location. The results show that average

sunburning UV has increased by up to a few percent per

decade between 1979 and 1992 over a wide range of latitudes (see Figure Q17-1). The largest increases are found

at high polar latitudes in both hemispheres. As expected,

the increases occur where decreases in total ozone are

observed to be the largest (see Figure Q13-1). The

smallest changes in sunburning UV are in the tropics,

where long-term total ozone changes are smallest.

UV Index changes. The “UV Index” is a measure of

daily surface UV levels that is relevant to the effects of

UV on human skin. The UV Index is used internationally

to increase public awareness about the detrimental effects

of UV on human health and to guide the need for protective measures. The Index is essentially a measure of erythemal radiation, of which UV-B is a principal component.

The daily maximum UV Index varies with location and

season, as shown for three locations in Figure Q17-2. The

highest daily values generally occur at the lowest latitudes

(tropics) and in summer when the midday Sun is closest to

overhead. Values in San Diego, California, for example,

normally are larger year round than those found in Barrow,

Alaska, which is at a higher latitude. At a given latitude,

UV Index values increase in mountainous regions. Winter

and fall values become zero in periods of continuous darkness at high-latitude locations. 

An illustrative example of how polar ozone depletion

increases the maximum daily UV Index is shown in

Figure Q17-2. Normal UV Index values for Palmer,

Antarctica, in spring were estimated from satellite measurements made during the period 1978-1983, before the

appearance of the “ozone hole” over Antarctica (see red

dotted line). In the last decade (1991-2001) severe and

persistent ozone depletion in spring has increased the UV

Index well above normal values for several months (see

thick red line). Now spring UV Index values in Palmer

(64°S) sometimes equal or exceed normal values measured in San Diego (32°N) in spring.

Other causes of long-term UV changes. The intensity of surface UV-B may also change as a result of other

human activities or climate change. Long-term changes

in cloudiness, aerosols, pollution, and snow or ice cover

will cause long-term changes in surface UV-B. At some

ground sites, measurements indicate that long-term

changes in UV-B have resulted from changes in one or

more of these factors. The impact of some of the changes

can be complex. For example, a change in cloud cover

usually results in a reduction of UV-B radiation below the

cloud layer, but can increase radiation above the layer (in

mountainous regions).

TWENTY QUESTIONS

Q.32

V. IMPLICATIONS OF OZONE DEPLETION

Q17: Does depletion of the ozone layer increase ground-level ultraviolet radiation?

Yes, ultraviolet radiation at Earth’s surface increases as the amount of overhead total ozone decreases, because

ozone absorbs ultraviolet radiation from the Sun. Measurements by ground-based instruments and estimates

made using satellite data have confirmed that surface ultraviolet radiation has increased in regions where

ozone depletion is observed.
UV changes and skin cancer. Skin cancer cases in

humans increase with the amount of sunburning UV

reaching Earth’s surface. Atmospheric scientists working

together with health professionals can estimate how skin

cancer cases will change in the future. The estimates are

based on knowing how sunburning UV increases as total

ozone is depleted and how total ozone depletion changes

with effective stratospheric chlorine (see Q16). Estimates

of future excess skin cancer cases are shown in Figure

Q15-1 using future estimates of effective chlorine based

on the 1992 Montreal Protocol provisions. The cases are

those that would occur in a population with the UV sensitivity and age distribution such as that of the United States.

The cases counted are those in excess of the number that

occurred in 1980 before ozone depletion was observed

(about 2000 per million population). The case estimates

include the fact that skin cancer in humans occurs long

after the exposure to sunburning UV. The results show

that, with current Protocol provisions, excess skin cancer

cases are predicted to increase in the early to middle

decades of the 21st century. By century’s end, with the

expected decreases in halogen source gas emissions, the

number of excess cases is predicted to return close to 1980

values. Without the provisions of the Protocol, excess

skin cancer cases would be expected to increase

unchecked throughout the century.

TWENTY QUESTIONS

Latitude

-80 -60 -40 -20 0 20 40 6 0

0

-4

4

8

12

16

Trend (percent per decade)

Sunburning UV

South North

Changes in Surface Ultraviolet Radiation

Average

Uncertainty range

0 8

Figure Q17-1. Changes in surface UV radiation.

Ultraviolet (UV) radiation at Earth’s surface that causes

sunburning has increased over the globe between 1979

and 1992. Also known as “erythemal radiation,” sunburning UV is harmful to humans and other life forms.

The increase shown here is estimated from observed

decreases in ozone and the relationship between ozone

and surface UV established at some surface locations.

The changes in ultraviolet radiation in the tropics are

estimated to be the smallest because observed changes

in tropical ozone have been the smallest.

Winter Spring Summer Fall 0

2

4

6

8

10

12

14

16

18

Maximum daily UV Index

1978-1983 1991-2001

Palmer, Antarctica (64°S)

San Diego, California (32°N)

Barrow, Alaska (71°N)

Seasonal Changes in the UV Index

20

Figure Q17-2. Changes in UV Index. The

maximum daily UV Index is a measure of

peak sunburning UV that occurs during the

day at a particular location. UV-B, which is

absorbed by ozone, is an important component of sunburning UV. The UV Index varies

with latitude and season as the Sun’s path

through the local sky changes. The highest

values of the maximum daily UV Index occur

in the tropics where the midday Sun is

highest throughout the year and where total

ozone values are lowest. For comparison,

the figure shows that the UV Index is higher

in San Diego than in Barrow throughout the

year. Index values are zero at high latitudes

when darkness is continuous. The effect of

ozone depletion on the Index is demonstrated by comparing the Palmer and San

Diego data in the figure. Normal values estimated for Palmer are shown for the 1978-

1983 period before the “ozone hole”

occurred each season (see red dotted line). In the last decade (1991-2001), Antarctic ozone depletion has led to an increase in

the maximum UV Index value at Palmer throughout spring (see yellow shaded region). Values at Palmer now sometimes equal

or exceed those measured in spring in San Diego, which is located at a much lower latitude. 

Q.33
Ozone depletion is not the major cause of climate

change, but ozone changes are linked to climate change

in important ways.

Radiative forcing of climate change. Human activities have led to the accumulation in the atmosphere of

several long-lived and radiatively active gases known as

“greenhouse gases.” Ozone is a greenhouse gas, along with

carbon dioxide (CO2), methane (CH4), nitrous oxide (N2O),

and halogen source gases. The accumulation of these gases

changes the radiative balance of Earth’s atmosphere. The

balance is between incoming solar radiation and outgoing

infrared radiation. Greenhouse gases generally change the

balance by absorbing outgoing radiation, leading to a

warming at Earth’s surface. This change in Earth’s radiative balance is called a radiative forcing of climate change.

If the total forcing from all gases becomes large enough,

climate could change in significant ways. A summary of

radiative forcings resulting from the increase in long-lived

greenhouse gases is shown in Figure Q18-1. All forcings

shown relate to human activities. Positive forcings lead to

warming and negative forcings lead to cooling of Earth’s

surface. The accumulation of carbon dioxide is the largest

forcing term. Carbon dioxide concentrations are increasing

in the atmosphere primarily as the result of the burning of

coal, oil, and natural gas for energy and transportation. The

atmospheric abundance of carbon dioxide is currently about

30% above what it was 150 years ago in preindustrial times. 

Stratospheric and tropospheric ozone. Changes in

stratospheric and tropospheric ozone both represent radiative forcings of climate change. Stratospheric ozone

absorbs solar radiation, which heats the stratosphere and

affects air motions and chemical reactions. Stratospheric

and tropospheric ozone both absorb infrared radiation

emitted by Earth’s surface, effectively trapping heat in the

atmosphere below. Overall, the depletion of stratospheric

ozone represents a negative radiative forcing. In contrast,

TWENTY QUESTIONS

Q.34

Q18: Is depletion of the ozone layer the principal cause of climate change?

No, ozone depletion itself is not the principal cause of climate change. However, because ozone is a greenhouse

gas, ozone changes and climate change are linked in important ways. Stratospheric ozone depletion and

increases in global tropospheric ozone that have occurred in recent decades both contribute to climate change.

These contributions to climate change are significant but small compared with the total contribution from all

other greenhouse gases. Ozone and climate change are indirectly linked because ozone-depleting gases, such

as the chlorofluorocarbons (CFCs), hydrochlorofluorocarbons (HCFCs), and halons, also contribute to climate

change.

Figure Q18-1. Climate change from

atmospheric gas changes. Human activities since 1750 have caused increases in the

abundances of several long-lived gases,

changing the radiative balance of Earth’s

atmosphere. These gases, known as

“greenhouse gases,” result in radiative forcings, which can lead to climate change. The

largest radiative forcings come from carbon

dioxide, followed by methane, tropospheric

ozone, the halogen-containing gases (see

Figure Q7-1), and nitrous oxide. Ozone

increases in the troposphere result from pollution associated with human activities. All

these forcings are positive, which leads to a

warming of Earth’s surface. In contrast,

stratospheric ozone depletion represents a

small negative forcing, which leads to

cooling of Earth’s surface. In the coming decades, halogen gas abundances and stratospheric ozone depletion are

expected to be reduced along with their associated radiative forcings. The link between these two forcing terms is an

important aspect of the radiative forcing of climate change. 

Decreases in

the stratosphere

3

2

1

0

Radiative Forcing of Climate Change from

Atmospheric Gas Changes (1750-2000)

Radiative forcing

(watts per square meter)

warming cooling

Carbon dioxide (CO2)

Methane (CH4)

Nitrous oxide (N2O)

Increases in

the troposphere

Increases in

long-lived gases

Ozone changes

-1

Halogen-containing gases
increases in tropospheric ozone due to surface pollution

represent a positive radiative forcing. The radiative

forcing due to tropospheric ozone increases is larger than

that associated with stratospheric ozone depletion. Both

forcing terms are significant, but are small in comparison

with the total forcing from all other greenhouse gases. 

Halogen source gases. An indirect link between

ozone depletion and climate change is the radiative

forcing from halogen source gases. These gases, which

include chlorofluorocarbons (CFCs), hydrochlorofluorocarbons (HCFCs), hydrofluorocarbons (HFCs), and

halons (see Figure Q7-1), are radiatively active in the

atmosphere before they are chemically converted in the

stratosphere. As a group, they represent a significant positive radiative forcing. Once converted, they form reactive halogen gases, which chemically destroy ozone. In

the coming decades, halogen gas abundances and their

associated positive radiative forcings are expected to

decrease (see Q16). With reduced halogen gases, stratospheric ozone depletion and its associated negative radiative forcing will also be reduced. This link between these

two forcing terms is an important aspect of the radiative

forcing of climate change. 

HFCs. The radiative forcing of halogen-containing

gases in Figure Q18-1 also includes that of the HFCs,

which do not cause ozone depletion. The HFCs are

increasing in the atmosphere because of their use as substitute gases for the CFCs and other halogen gases. HFCs

are not regulated under the Montreal Protocol, but are

included in the group of gases listed in the Kyoto Protocol

of the United Nations Framework Convention on Climate

Change (UNFCC). 

Climate change. Certain changes in Earth’s climate

could affect the future of the ozone layer. Stratospheric

ozone is influenced by changes in temperatures and winds

in the stratosphere. For example, low temperatures and

strong polar winds both affect the extent and severity of

winter polar ozone depletion. While Earth’s surface is

expected to warm in response to the positive radiative

forcing from carbon dioxide increases, the stratosphere

is expected to cool. Indeed, a small cooling of the lower

stratosphere has occurred since the 1970s. A cooler stratosphere would extend the time period over which polar

stratospheric clouds (PSCs) are present in polar regions

and, as a result, might increase winter ozone depletion.

These changes could delay the recovery of the ozone

layer. In the upper stratosphere at altitudes above PSC

formation regions, a cooler stratosphere is expected to

increase ozone amounts. Changes in atmospheric composition that lead to a warmer climate may also affect the

balance of production and loss processes of stratospheric

ozone (see Q20).

TWENTY QUESTIONS

Q.35
Recovery factors. Detecting the recovery of the

ozone layer will rely on comparisons of the latest ozone

values with values measured in the past. Because of its

importance, ozone likely will be measured continuously

in the future using a variety of techniques and measurement platforms (see Q5). In the ozone comparisons, scientists will look for improvements in certain factors

related to the distribution of ozone. These factors include

the following:

• A lessening of the decline in global ozone, either

in total ozone or ozone at specific altitudes in the

stratosphere. 

• An increase in global total ozone amounts toward

values found before 1980 when halogen source gas

abundances in the atmosphere were much lower

than today.

• A sustained reduction in the maximum size of the

Antarctic “ozone hole.” 

• A sustained increase in the minimum value of

ozone found in the Antarctic ozone hole. 

• Less ozone depletion in Arctic winters during

which temperatures are below polar stratospheric

cloud (PSC) formation temperatures. 

As the ozone layer approaches full recovery, scientists expect to observe improvements in all these factors. 

Natural factors. Global total ozone is influenced by

two important natural factors, namely, changes in the output

of the Sun and volcanic eruptions (see Q14). The evaluation of ozone recovery must include the effects of these natural factors. The solar effect on ozone is expected to be

predictable based on the well-established 11-year cycle of

solar output. Volcanic eruptions are particularly important

because they increase ozone depletion caused by reactive

halogen gases, but cannot be predicted. The occurrence of

a large volcanic eruption in the next decades when effective stratospheric chlorine levels are still high (see Figure

Q16-1) may obscure progress in overall ozone recovery by

temporarily increasing ozone depletion. The natural variation of ozone amounts will also limit how easily small

improvements in ozone abundances can be detected.

TWENTY QUESTIONS

Q.36

VI. STRATOSPHERIC OZONE IN THE FUTURE

Q19: How will recovery of the ozone layer be detected?

Scientists expect to detect the recovery of the ozone layer with careful comparisons of the latest ozone measurements with past values. Changes in total overhead ozone at various locations and in the extent and severity of

the Antarctic “ozone hole” will be important factors in gauging ozone recovery. Natural variations in ozone

amounts will limit how soon recovery can be detected with future ozone measurements. 
Halogen source gas reductions. Ozone depletion

caused by human-produced chlorine and bromine gases is

expected to gradually disappear by about the middle of

the 21st century as the abundances of these gases decline

in the stratosphere. The decline will follow the reductions

in emissions that are expected to occur under the provisions of the Montreal Protocol and its Adjustments and

Amendments (see Figure Q16-1). The emission reductions are based on the assumption of full compliance by

the developed and developing nations of the world. The

slowing of the increases in the atmospheric abundances

of several halogen gases and the substantial reduction of

one principal halogen gas, methyl chloroform, has already

occurred. Natural chemical and transport processes limit

the rate at which halogen gases can be removed from the

stratosphere. The atmospheric chemical lifetimes of the

halogen source gases range up to 100 years (see Table Q7-1).

Chlorofluorocarbon-12 (CFC-12), with a lifetime of 100

years, will require about 200 to 300 years before it is

removed (less than 5% remaining) from the atmosphere

(see Figure Q16-1). 

Ozone predictions. Computer models of the atmosphere are used to assess past changes in global ozone and

predict future changes. Two important measures of ozone

considered by scientists are global total ozone averaged

between 60°N and 60°S latitudes, and minimum ozone

values in the Antarctic “ozone hole.” Both measures show

ongoing ozone depletion that began in the 1980s (see

Figure Q20-1). The range of model predictions shows

that the lowest ozone values are expected to occur before

2020 and that substantial recovery of ozone is expected

by the middle of the 21st century. The range of predicTWENTY QUESTIONS

Q.37

Year

1980

-6

Recovery of Global Ozone

2000 2020 2040

-4

-2

0

2

4

60°N- 60°S total ozone

Range of atmospheric

model predictions

Observations:

Average and range

Antarctic total ozone

Range of atmospheric

model predictions

Satellite observations

200

100

0

Minimum total ozone (DU) Ozone change from 1980 (%)

Figure Q20-1. Global ozone recovery predictions.

Observed values of global total ozone (top panel) and minimum total ozone over Antarctica (bottom panel) have

decreased beginning in the early 1980s. As halogen source

gas emissions decrease in the early 21st century, ozone

values are expected to increase and recover toward pre1980 values. Atmospheric computer models that account

for changes in halogen gases and other atmospheric parameters can be used to predict how ozone amounts will

increase. These model results show that recovery is

expected to be significant by 2050, or perhaps earlier. The

range of model predictions comes from the use of several

different models that have different assumptions about the

future climate and composition of the atmosphere. These

assumptions are an attempt to account for estimated differences in atmospheric composition and other parameters

between 1980, before the “ozone hole” began, and 2050.

Q20: When is the ozone layer expected to recover?

The ozone layer is expected to recover by the middle of the 21st century, assuming global compliance with the

Montreal Protocol. Chlorine- and bromine-containing gases that cause ozone depletion will decrease in the

coming decades under the provisions of the Protocol. However, volcanic eruptions in the next decades could

delay ozone recovery by several years and the influence of climate change could accelerate or delay ozone

recovery.
tions comes from several computer models of the atmosphere that have different assumptions concerning the

future climate and composition of the atmosphere. Some

of these models indicate that recovery of total ozone may

possibly come well before midcentury.

A different atmosphere in 2050. By the middle of

the 21st century, halogen amounts in the stratosphere are

expected to be similar to those present in 1980 before the

onset of the ozone hole (see Figure Q16-1). However,

other aspects of the global atmosphere will not be the same

in 2050 as in 1980. The ozone recovery evaluations in

Figure Q20-1 attempt to take these differences into

account. For example, since 1980 human activities have

increased the abundance of important greenhouse gases,

including carbon dioxide, methane, and nitrous oxide.

The accumulation of these gases is expected to cause

warmer surface temperatures and colder stratospheric

temperatures. Colder temperatures may accelerate ozone

recovery in the upper stratosphere (about 40 kilometers

(25 miles) altitude). Colder winter temperatures over

polar regions will increase the occurrences of polar stratospheric clouds (PSCs) and chemical ozone destruction

(see Q10). Water vapor increases that have occurred in

the stratosphere over the last two decades also will lead to

increased PSC occurrences and associated ozone destruction. A cooler, wetter polar stratosphere could delay ozone

recovery beyond what would be predicted for the 1980

atmosphere. Increased methane and nitrous oxide abundances due to human activities also cause some change in

the overall balance of the chemical production and

destruction of global stratospheric ozone. Finally, an outcome not included in models is the occurrence of one or

more large volcanic eruptions in the coming decades. In

that case, stratospheric particles may increase for several

years, temporarily reducing global ozone amounts (see

Q14) and delaying the recovery of the ozone layer. 

TWENTY QUESTIONS

Q.38
APPENDIX A

LIST OF INTERNATIONAL AUTHORS,

CONTRIBUTORS, AND REVIEWERS

COCHAIRS

Ayité-Lô Nohende Ajavon Université de Lomé Togo

Daniel L. Albritton National Oceanic and Atmospheric Administration US

Gérard Mégie Centre National de la Recherche Scientifique France

Robert T. Watson World Bank US

AUTHORS AND CONTRIBUTORS

CHAPTER 1 

CONTROLLED SUBSTANCES AND OTHER SOURCE GASES

Chapter Lead Authors

Stephen A. Montzka NOAA Climate Monitoring and Diagnostics Laboratory US

Paul J. Fraser CSIRO Division of Atmospheric Research Australia

Coauthors

James H. Butler NOAA Climate Monitoring and Diagnostics Laboratory US

Peter S. Connell Lawrence Livermore National Laboratory US

Derek M. Cunnold Georgia Institute of Technology US

John S. Daniel NOAA Aeronomy Laboratory US

Richard G. Derwent UK Meteorological Office UK

Shyam Lal Physical Research Laboratory India

Archie McCulloch Marbury Technical Consulting UK

David E. Oram University of East Anglia UK

Claire E. Reeves University of East Anglia UK

Eugenio Sanhueza Instituto Venezolano de Investigaciones Cientificas Venezuela

L. Paul Steele CSIRO Division of Atmospheric Research Australia

Guus J.M. Velders National Institute of Public Health and the Environment (RIVM) The Netherlands

Ray F. Weiss University of California at San Diego/Scripps Institution US

of Oceanography

Rodolphe J. Zander Université de Liège Belgium

Contributors

Stephen O. Andersen Environmental Protection Agency US

James Anderson Harvard University US

Donald R. Blake University of California at Irvine US

A.1
Martyn P. Chipperfield University of Leeds UK

Ed Dlugokencky NOAA Climate Monitoring and Diagnostics Laboratory US

James W. Elkins NOAA Climate Monitoring and Diagnostics Laboratory US

Andreas Engel Universität Frankfurt Germany

David B. Harper The Queen's University of Belfast UK

Emmanuel Mahieu Université de Liège Belgium

Klaus Pfeilsticker Universität Heidelberg Germany

Jean-Pierre Pommereau Service d'Aéronomie du CNRS France

James M. Russell III Hampton University US

Gary Taylor Taylor/Wagner, Inc. Canada

Michel Van Roozendael Institut d'Aeronomie Spatiale de Belgique Belgium

Darryn W. Waugh Johns Hopkins University US

CHAPTER 2

VERY SHORT-LIVED HALOGEN AND SULFUR SUBSTANCES

Chapter Lead Authors

Malcolm K.W. Ko Atmospheric and Environmental Research, Inc. US

Gilles Poulet CNRS-Université d'Orléans France

Coauthors

Donald R. Blake University of California at Irvine US

Olivier Boucher CNRS-Université des Sciences et Technologies de Lille France

James H. Burkholder NOAA Aeronomy Laboratory US

Mian Chin Georgia Institute of Technology/NASA Goddard Space Flight Center US

R. Anthony Cox University of Cambridge UK

Christian George Laboratoire d'Application de la Chimie à l'Environnement/ France

University of Lyon

Hans-F. Graf Max-Planck-Institut für Meteorologie Germany

James R. Holton University of Washington US

Daniel J. Jacob Harvard University US

Katherine S. Law University of Cambridge UK

Mark G. Lawrence Max-Planck-Institut für Chemie Germany

Pauline M. Midgley M&D Consulting Germany

Paul W. Seakins University of Leeds UK

Dudley E. Shallcross University of Bristol UK

Susan E. Strahan University of Maryland US

Donald J. Wuebbles University of Illinois US

Yoko Yokouchi National Institute for Environmental Studies Japan

Contributors

Nicola J. Blake University of California at Irvine US

James H. Butler NOAA Climate Monitoring and Diagnostics Laboratory US

Anne R. Douglass NASA Goddard Space Flight Center US

Victor L. Dvortsov PacifiCorp US

Ian Folkins Dalhousie University Canada

Peter H. Haynes University of Cambridge UK

Abdelwahid Mellouki CNRS-Laboratoire de Combustion et Systèmes Réactifs France

Michael J. Prather University of California at Irvine US

AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.2
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.3

José M. Rodríguez University of Miami US

Sue M. Schauffler National Center for Atmospheric Research US

Theodore G. Shepherd University of Toronto Canada

Christiane Textor Max-Planck-Institut für Meteorologie Germany

Claudia Timmreck Max-Planck-Institut für Meteorologie Germany

Debra K. Weisenstein Atmospheric and Environmental Research, Inc. US

CHAPTER 3

POLAR STRATOSPHERIC OZONE: PAST AND FUTURE

Chapter Lead Authors

Paul A. Newman NASA Goddard Space Flight Center US

John A. Pyle Centre for Atmospheric Science, University of Cambridge UK

Coauthors

John Austin UK Meteorological Office UK

Geir O. Braathen Norwegian Institute for Air Research (NILU) Norway

Pablo O. Canziani CONICET/Universidad de Buenos Aires Argentina

Ken S. Carslaw University of Leeds UK

Piers M. de F. Forster University of Reading UK

Sophie Godin-Beekmann Service d'Aeronomie du CNRS France

Bjørn M. Knudsen Danmarks Meteorologiske Institut Denmark

Karin Kreher National Institute of Water and Atmospheric Research (NIWA) New Zealand

Hideaki Nakane National Institute for Environmental Studies Japan

Steven Pawson Goddard Earth Sciences and Technology Center/University of Maryland US

V. Ramaswamy NOAA Geophysical Fluid Dynamics Laboratory US

Markus Rex Alfred Wegener Institute for Polar and Marine Research Germany

Ross J. Salawitch California Institute of Technology/NASA Jet Propulsion Laboratory US

Drew T. Shindell NASA Goddard Institute for Space Studies US

Azadeh Tabazadeh NASA Ames Research Center US

Darin W. Toohey University of Colorado US

Contributors

Douglas R. Allen Naval Research Laboratory US

Linnea M. Avallone University of Colorado US

Stephen R. Beagley York University Canada

Greg E. Bodeker National Institute of Water and Atmospheric Research (NIWA) New Zealand

Christoph Brühl Max-Planck-Institut für Chemie Germany

John Christy University of Alabama at Huntsville US

Brian Connor National Institute of Water and Atmospheric Research (NIWA) New Zealand

Martin Dameris DLR Institut für Physik der Atmosphäre Germany

Anne R. Douglass NASA Goddard Space Flight Center US

Stephen Eckermann Naval Research Laboratory US

Melvyn Gelman NOAA NWS Climate Prediction Center US

Florence Goutail Service d'Aeronomie du CNRS France

Patrick Hamill San Jose State University US

Yuri Koshelkov Central Aerological Observatory Russia

Karin Labitzke Freie Universität Berlin Germany

Ulrike Langematz Freie Universität Berlin Germany
Roger Lin NOAA National Centers for Environmental Prediction/ US

RS Information Systems

Elisa Manzini Max-Planck-Institut für Meteorologie Germany

Tatsuya Nagashima National Institute for Environmental Studies Japan

Eric R. Nash Science Systems and Applications, Inc. US

John Nash UK Meteorological Office UK

Samuel J. Oltmans NOAA Climate Monitoring and Diagnostics Laboratory US

David Parker Hadley Centre Met Office UK

Klaus Pfeilsticker Universität Heidelberg Germany

Giovanni Pitari Università L'Aquila Italy

William J. Randel National Center for Atmospheric Research US

Eugene Rozanov World Radiation Center and Institute for Atmospheric and Switzerland

Climate Science ETH

Michelle L. Santee NASA Jet Propulsion Laboratory US

Christina Schnadt DLR Institut für Physik der Atmosphäre Germany

M. Daniel Schwarzkopf NOAA Geophysical Fluid Dynamics Laboratory US

Theodore G. Shepherd University of Toronto Canada

Masanori Shitamichi Japan Meteorological Agency Japan

Peter von der Gathen Alfred Wegener Institute for Polar and Marine Research Germany

Darryn W. Waugh Johns Hopkins University US

Paul O. Wennberg California Institute of Technology US

CHAPTER 4

GLOBAL OZONE: PAST AND FUTURE

Chapter Lead Authors

Martyn P. Chipperfield University of Leeds UK

William J. Randel National Center for Atmospheric Research US

Coauthors

Greg E. Bodeker National Institute of Water and Atmospheric Research (NIWA) New Zealand

Martin Dameris DLR Institut für Physik der Atmosphäre Germany

Vitali E. Fioletov Meteorological Service of Canada Canada

Randall R. Friedl NASA Jet Propulsion Laboratory US

Neil R.P. Harris European Ozone Research Coordinating Unit UK

Jennifer A. Logan Harvard University US

Richard D. McPeters NASA Goddard Space Flight Center US

Nzioka John Muthama University of Nairobi Kenya

Thomas Peter Institute for Atmospheric and Climate Science ETH-Zurich Switzerland

Theodore G. Shepherd University of Toronto Canada

Keith P. Shine University of Reading UK

Susan Solomon NOAA Aeronomy Laboratory US

Larry W. Thomason NASA Langley Research Center US

Joseph M. Zawodny NASA Langley Research Center US

Contributors

John Austin UK Meteorological Office UK

Michel Bourqui University of Reading UK

Peter Braesicke University of Cambridge UK

AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.4
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.5

Christoph Brühl Max-Planck-Institut für Chemie Germany

Neal Butchart UK Meteorological Office UK

David B. Considine NASA Langley Research Center US

Derek M. Cunnold Georgia Institute of Technology US

David W. Fahey NOAA Aeronomy Laboratory US

Eric L. Fleming NASA Goddard Space Flight Center US

Marvin A. Geller State University of New York, Stony Brook US

Sophie Godin-Beekmann Service d'Aeronomie du CNRS France

Volker Grewe DLR Institut für Physik der Atmosphäre Germany

Joanna D. Haigh Imperial College of Science, Technology, and Medicine UK

Charles H. Jackman NASA Goddard Space Flight Center US

Paul Johnston National Institute of Water and Atmospheric Research (NIWA) New Zealand

Ulrike Langematz Freie Universität Berlin Germany

Katherine S. Law University of Cambridge UK

J. Ben Liley National Institute of Water and Atmospheric Research (NIWA) New Zealand

Inna A. Megretskaia Harvard University US

Alvin J. Miller NOAA NWS Climate Prediction Center US

Tatsuya Nagashima National Institute for Environmental Studies Japan

Giovanni Pitari Università L'Aquila Italy

Robert W. Portmann NOAA Aeronomy Laboratory US

V. Ramaswamy NOAA Geophysical Fluid Dynamics Laboratory US

Bjørg Rognerud Universitetet I Oslo Norway

Joan E. Rosenfield NASA Goddard Space Flight Center US

Martin N. Ross Aerospace Corporation US

Christina Schnadt DLR Institut für Physik der Atmosphäre Germany

M. Daniel Schwarzkopf NOAA Geophysical Fluid Dynamics Laboratory US

Drew T. Shindell NASA Goddard Institute for Space Studies US

Claire A. Smith Imperial College of Science, Technology, and Medicine UK

Sergei Smyshlyaev Russian State Hydrometeorological University/ Russia

State University of New York

Richard S. Stolarski NASA Goddard Space Flight Center US

Guus J.M. Velders National Institute of Public Health and the Environment (RIVM) The Netherlands

Ray Wang Georgia Institute of Technology US

Debra K. Weisenstein Atmospheric and Environmental Research, Inc. US

Fei Wu National Center for Atmospheric Research US

Rodolphe J. Zander Université de Liège Belgium

CHAPTER 5

SURFACE ULTRAVIOLET RADIATION: PAST AND FUTURE

Chapter Lead Authors

James B. Kerr Meteorological Service of Canada Canada

Gunther Seckmeyer Universität Hannover Germany

Coauthors

Alkiviadis F. Bais Aristotle University of Thessaloniki Greece

Germar Bernhard Biospherical Instruments, Inc. US

Mario Blumthaler Universität Innsbruck Austria

Susana B. Diaz Centro Austral de Investigaciones Cientificas (CADIC) Argentina

Nickolay A. Krotkov University of Maryland US
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.6

Daniel Lubin Scripps Institution of Oceanography US

Richard L. McKenzie National Institute of Water and Atmospheric Research (NIWA) New Zealand

Ali A. Sabzipavar University of Bou-Ali Sina Iran

Jean Verdebout European Commission Joint Research Centre Italy

Contributors

Antti Arola Finnish Meteorological Institute Finland

Vitali E. Fioletov Meteorological Service of Canada Canada

Jay R. Herman NASA Goddard Space Flight Center US

Peter W. Kiedron State University of New York at Albany US

J. Ben Liley National Institute of Water and Atmospheric Research (NIWA) New Zealand

Craig S. Long NOAA NWS Climate Prediction Center US

Sasha Madronich National Center for Atmospheric Research US

Timothy J. Martin Universität Graz Austria

Ralf Meerkötter DLR-Institut für Physik der Atmosphäre Germany

Patrick J. Neale Smithsonian Environmental Research Center US

Harry Slaper National Institute of Public Health and the Environment (RIVM) The Netherlands

James R. Slusser Colorado State University US

Petteri Taalas Finnish Meteorological Institute Finland

Michiel van Weele Koninklijk Nederlands Meteorologisch Instituut (KNMI) The Netherlands

Elizabeth C. Weatherhead NOAA Air Resources Laboratory US

Christos S. Zerefos Aristotle University of Thessaloniki Greece

TWENTY QUESTIONS AND ANSWERS ABOUT THE OZONE LAYER

Lead Author

David W. Fahey NOAA Aeronomy Laboratory US

CHAPTER EDITORIAL CONTRIBUTORS

CHAPTER 1: CONTROLLED SUBSTANCES AND OTHER SOURCE GASES

Nada Derek CSIRO Division of Atmospheric Research Australia

CHAPTER 3: POLAR STRATOSPHERIC OZONE: PAST AND FUTURE

Rose Kendall Computer Sciences Corporation US

Kathy A. Thompson Computer Sciences Corporation US

REVIEWERS

Alberto Adriani Consiglio Nazionale delle Ricerche Istituto di Fisica dell'Atmosfera Italy

Ayité-Lô Nohende Ajavon Université de Lomé Togo

Daniel L. Albritton NOAA Aeronomy Laboratory US

Georgios T. Amanatidis European Commission Belgium

Stephen O. Andersen Environmental Protection Agency US
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.7

Gustavo A. Argüello Universidad Nacional de Córdoba Argentina

Roger Atkinson University of California at Riverside US

Pieter J. Aucamp Ptersa Environmental Consultant South Africa

John Austin UK Meteorological Office UK

Mark P. Baldwin NorthWest Research Associates, Inc. US

Pranvera Bekteshi Hydrometeorological Institute Albania

Germar Bernhard Biospherical Instruments, Inc. US

Donald R. Blake University of California at Irvine US

Greg E. Bodeker National Institute of Water and Atmospheric Research (NIWA) New Zealand

Rumen D. Bojkov Centre for International Postgraduate Studies of Germany

Environmental Management 

Guy P. Brasseur Max-Planck-Institut für Meteorologie Germany

Bram Bregman Royal Netherlands Meteorological Institute The Netherlands

Christoph Brühl Max-Planck-Institut für Chemie Germany

William Brune Pennsylvania State University US

James H. Butler NOAA Climate Monitoring and Diagnostics Laboratory US

James Calm Engineering Consultant US

Ken S. Carslaw University of Leeds UK

Marie-Lise Chanin Service d’Aeronomie du CNRS France

Martyn P. Chipperfield University of Leeds UK

R. Anthony Cox University of Cambridge UK

Derek M. Cunnold Georgia Institute of Technology US

Martin Dameris DLR Institut für Physik der Atmosphäre Germany

John S. Daniel NOAA Aeronomy Laboratory US

Susana B. Diaz Centro Austral de Investigaciones Cientificas (CADIC) Argentina

Katja Drdla NASA Ames Research Center US

Ellsworth S. Dutton NOAA Climate Monitoring and Diagnostics Laboratory US

Thomas Duafala Tri-Cal Research Division US

Ezequiel Echer Instituto Nacional de Pesquisas Espaciais (INPE) Brazil

Kalju Eerme Tartu Observatory Estonia

James W. Elkins NOAA Climate Monitoring and Diagnostics Laboratory US

Andreas Engel Universität Frankfurt Germany

Christine A. Ennis NOAA Aeronomy Laboratory/CIRES US

David W. Fahey NOAA Aeronomy Laboratory US

Joe Farman European Ozone Research Coordinating Unit UK

Vitali E. Fioletov Meteorological Service of Canada Canada

Piers M. de F. Forster University of Reading UK

Paul J. Fraser CSIRO Division of Atmospheric Research Australia

Ru-Shan Gao NOAA Aeronomy Laboratory US

Rolando R. Garcia National Center for Atmospheric Research US

Marvin A. Geller State University of New York, Stony Brook US

Sophie Godin-Beekmann Service d'Aeronomie du CNRS France

Marco González United Nations Environment Programme Kenya

Michael Graber United Nations Environment Programme Kenya

Hans-F. Graf Max-Planck-Institut für Meteorologie Germany

Joanna D. Haigh Imperial College of Science, Technology, and Medicine UK

Neil R.P. Harris European Ozone Research Coordinating Unit UK

Didier Hauglustaine CNRS-Laboratoire des Sciences du Climat et de l'Environnement France

Peter H. Haynes University of Cambridge UK
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.8

Jay R. Herman NASA Goddard Space Flight Center US

David J. Hofmann NOAA Climate Monitoring and Diagnostics Laboratory US

James R. Holton University of Washington US

Robert D. Hudson University of Maryland US

Drusilla Hufford Environmental Protection Agency US

Abdelmoneim A. Ibrahim Egyptian Meteorological Authority Egypt

Mohammad Ilyas University of Science Malaysia Malaysia

Takashi Imamura National Institute for Environmental Studies Japan

Ivar S.A. Isaksen University of Oslo Norway

Charles H. Jackman NASA Goddard Space Flight Center US

Daniel J. Jacob Harvard University US

Mauricio Jaramillo-Ayerbe Pontificia Universidad Javeriana-Cali Colombia

David Karoly Monash University Australia

Nozomi Kawamoto National Space Development Agency Japan

Jack A. Kaye NASA Office of Earth Science US

James B. Kerr Meteorological Service of Canada Canada

M.A.K. Khalil Portland State University US

Dieter Kley Institut für Chemie und Dynamik der Geoshäre Forschungszentrum Jülich Germany

Malcolm K.W. Ko Atmospheric and Environmental Research, Inc. US

Yutaka Kondo University of Tokyo Japan

Janusz W. Krzys ́cin Polish Academy of Sciences Poland

Lambert Kuijpers Technical University Pav The Netherlands

Michael J. Kurylo NASA Headquarters US

Murari Lal Indian Institute of Technology India

Shyam Lal Physical Research Laboratory India

Kathleen O. Lantz NOAA Air Resources Laboratory US

Neils Larsen Danish Meteorological Institute Denmark

Katherine S. Law University of Cambridge UK

Mark G. Lawrence Max-Planck-Institut für Chemie Germany

J. Ben Liley National Institute of Water and Atmospheric Research (NIWA) New Zealand

Shaw Liu Academia Sinica Taiwan R.O.C.

Jennifer A. Logan Harvard University US

A. Robert MacKenzie Lancaster University UK

Gloria L. Manney NASA Jet Propulsion Laboratory/New Mexico Highlands University US

Céline Mari Laboratoire d'Aérologie, Observatoire Midi Pyrénées France

W. Andrew Matthews National Institute of Water and Atmospheric Research (NIWA) New Zealand

Konrad Mauersberger Max-Planck-Institut für Kernphysik Germany

Archie McCulloch Marbury Technical Consulting UK

Gordon McFadyen Scottish Environment Protection Agency UK

Mack McFarland E.I. DuPont de Nemours & Company US

Daniel S. McKenna National Center for Atmospheric Research US

Richard L. McKenzie National Institute of Water and Atmospheric Research (NIWA) New Zealand

Ralf Meerkötter DLR-Institut für Physik der Atmosphäre Germany

Gérard Mégie Centre National de la Recherche Scientifique France

Davit Melkonyan Department of Hydrometeorology Armenia

Pauline M. Midgley M&D Consulting Germany

Stephen A. Montzka NOAA Climate Monitoring and Diagnostics Laboratory US

Rolf Müller Forschungszentrum Jülich GmbH Germany

Nzioka John Muthama University of Nairobi Kenya
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.9

Hideaki Nakane National Institute for Environmental Studies Japan

Eric R. Nash Science Systems and Applications, Inc. US

Paul A. Newman NASA Goddard Space Flight Center US

Samuel J. Oltmans NOAA Climate Monitoring and Diagnostics Laboratory US

Alan O'Neill University of Reading UK

Michael Oppenheimer Princeton University US

Eduardo Palenque Instituto de Investigaciones Físcias, Universidad Mayor de San Andres Bolivia

Panos Papagiannakopuolos University of Crete Greece

Steven Pawson Goddard Earth Sciences and Technology Center/University of Maryland US

Stuart A. Penkett University of East Anglia UK

Sunil Kumar Peshin India Meteorological Department India

Thomas Peter Institute for Atmospheric and Climate Science ETH-Zurich Switzerland

Klaus Pfeilsticker Universität Heidelberg Germany

Giovanni Pitari Università L'Aquila Italy

Ulrich Platt University of Heidelberg Germany

Ian Plumb CSIRO Telecommunications and Industrial Physics Australia

Jean-Pierre Pommereau Service d'Aéronomie du CNRS France

Lamont R. Poole NASA Langley Research Center US

Gilles Poulet CNRS-Université d'Orléans France

Michael J. Prather University of California at Irvine US

Margarita Préndez Universidad de Chile Chile

Ronald Prinn Massachusetts Institute of Technology US

Michael H. Proffitt World Meteorological Organization Switzerland

John A. Pyle Centre for Atmospheric Science, University of Cambridge UK

S. Ramachandran Physical Research Laboratory India

V. Ramaswamy NOAA Geophysical Fluid Dynamics Laboratory US

William J. Randel National Center for Atmospheric Research US

Lakshman Randeniya CSIRO Telecommunications and Industrial Physics Australia

Philip J. Rasch National Center for Atmospheric Research US

A.R. Ravishankara NOAA Aeronomy Laboratory US

Claire E. Reeves University of East Anglia UK

Brian A. Ridley National Center for Atmospheric Research US

Curtis P. Rinsland NASA Langley Research Center US

Henning Rodhe University of Stockholm Sweden

José M. Rodríguez University of Miami US

Martin N. Ross Aerospace Corporation US

Nelson A. Sabogal United Nations Environment Programme Kenya

Ross J. Salawitch California Institute of Technology/NASA Jet Propulsion Laboratory US

Eugenio Sanhueza Instituto Venezolano de Investigaciones Cientificas Venezuela

Michelle L. Santee NASA Jet Propulsion Laboratory US

Toru Sasaki Japan Meteorological Agency Japan

Yasuhiro Sasano National Institute for Environmental Studies Japan

Sue M. Schauffler National Center for Atmospheric Research US

Ulrich Schmidt Universität of Frankfurt Germany

Ulrich Schumann DLR Institut für Physik der Atmosphäre Germany

Gunther Seckmeyer Universität Hannover Germany

Dian J. Seidel NOAA Air Resources Laboratory US

Theodore G. Shepherd University of Toronto Canada

Keith P. Shine University of Reading UK
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.10

Peter G. Simmonds University of Bristol UK

Paul C. Simon Institut d'Aeronomie Spatiale de Belgique Belgium

James R. Slusser Colorado State University US

Susan Solomon NOAA Aeronomy Laboratory US

Johannes Stähelin Swiss Federal Institute of Technology-Zurich Switzerland

Andreas Stohl Technical University Munich Germany

B.H. Subbaraya Indian Space Research Organization Headquarters India

Christine C. Sweet Science and Technology Corporation US

Petteri Taalas Finnish Meteorological Institute Finland

David W. Tarasick Meteorological Service of Canada Canada

Manfred Tevini Botanisches Institut Denmark

David Thompson Colorado State University US

Geoffrey C. Toon NASA Jet Propulsion Laboratory US

Adrian F. Tuck NOAA Aeronomy Laboratory US

Jan C. van der Leun Ecofys The Netherlands

Guus J.M. Velders National Institute of Public Health and the Environment (RIVM) The Netherlands

Jean Verdebout European Commission Joint Research Centre Italy

Mario Visca Ausimont S.p.A. Italy

C. Michael Volk Universität Frankfurt Germany

Joe W. Waters NASA Jet Propulsion Laboratory US

Robert T. Watson World Bank US

Elizabeth C. Weatherhead NOAA Air Resources Laboratory US

Philipp Weihs Universität für Bodenkultur Austria

Ray F. Weiss University of California at San Diego/Scripps Institution of Oceanography US

Paul O. Wennberg California Institute of Technology US

Donald J. Wuebbles University of Illinois US

Vladimir Yushkov Central Aerological Observatory Russia

Rodolphe J. Zander Université de Liège Belgium

Francesco Zaratti Instituto de Investigaciones Físicas, Universidad Mayor de San Andres Bolivia

Christos S. Zerefos Aristotle University of Thessaloniki Greece

Xiuji Zhou Chinese Academy of Meteorological Sciences China

Tong Zhu Peking University China

OZONE PEER-REVIEW MEETING

Les Diablerets, Switzerland

24-28 June 2002

Ayité-Lô Nohende Ajavon Université de Lomé Togo

Daniel L. Albritton NOAA Aeronomy Laboratory US

Georgios T. Amanatidis European Commission Belgium

Stephen O. Andersen Environmental Protection Agency US

Pieter J. Aucamp Ptersa Environmental Consultant South Africa

John Austin UK Meteorological Office UK

Germar Bernhard Biospherical Instruments, Inc. US

Greg E. Bodeker National Institute of Water and Atmospheric Research (NIWA) New Zealand

James H. Butler NOAA Climate Monitoring and Diagnostics Laboratory US
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.11

Ken S. Carslaw University of Leeds UK

R. Anthony Cox University of Cambridge UK

Derek M. Cunnold Georgia Institute of Technology US

Martin Dameris DLR Institut für Physik der Atmosphäre Germany

Susana B. Diaz Centro Austral de Investigaciones Cientificas (CADIC) Argentina

Thomas Duafala Tri-Cal Research Division US

Christine A. Ennis NOAA Aeronomy Laboratory/CIRES US

David W. Fahey NOAA Aeronomy Laboratory US

Vitali E. Fioletov Meteorological Service of Canada Canada

Rolando R. Garcia National Center for Atmospheric Research US

Sophie Godin-Beekmann Service d'Aeronomie du CNRS France

Michael Graber United Nations Environment Programme Kenya

Neil R.P. Harris European Ozone Research Coordinating Unit UK

Peter H. Haynes University of Cambridge UK

Abdelmoneim A. Ibrahim Egyptian Meteorological Authority Egypt

Mohammad Ilyas University of Science Malaysia Malaysia

Nozomi Kawamoto National Space Development Agency Japan

James B. Kerr Meteorological Service of Canada Canada

Malcolm K.W. Ko Atmospheric and Environmental Research, Inc. US

Michael J. Kurylo NASA Headquarters US

Neils Larsen Danish Meteorological Institute Denmark

Katherine S. Law University of Cambridge UK

Konrad Mauersberger Max-Planck-Institut für Kernphysik Germany

Archie McCulloch Marbury Technical Consulting UK

Mack McFarland E.I. DuPont de Nemours & Company US

Gérard Mégie Centre National de la Recherche Scientifique France

Pauline M. Midgley M&D Consulting Germany

Stephen A. Montzka NOAA Climate Monitoring and Diagnostics Laboratory US

Rolf Müller Forschungszentrum Jülich GmbH Germany

Nzioka John Muthama University of Nairobi Kenya

Hideaki Nakane National Institute for Environmental Studies Japan

Paul A. Newman NASA Goddard Space Flight Center US

Steven Pawson Goddard Earth Sciences and Technology Center/University of Maryland US

Stuart A. Penkett University of East Anglia UK

Ian Plumb CSIRO Telecommunications and Industrial Physics Australia

Jean-Pierre Pommereau Service d'Aéronomie du CNRS France

Lamont R. Poole NASA Langley Research Center US

Gilles Poulet CNRS-Université d'Orléans France

Margarita Préndez Universidad de Chile Chile

Michael H. Proffitt World Meteorological Organization Switzerland

John A. Pyle Centre for Atmospheric Science, University of Cambridge UK

V. Ramaswamy NOAA Geophysical Fluid Dynamics Laboratory US

William J. Randel National Center for Atmospheric Research US

A.R. Ravishankara NOAA Aeronomy Laboratory US

Claire E. Reeves University of East Anglia UK

José M. Rodríguez University of Miami US

Nelson A. Sabogal United Nations Environment Programme Kenya

Ross J. Salawitch California Institute of Technology/NASA Jet Propulsion Laboratory US

Gunther Seckmeyer Universität Hannover Germany
A.12

Dian J. Seidel NOAA Air Resources Laboratory US

Theodore G. Shepherd University of Toronto Canada

Keith P. Shine University of Reading UK

Paul C. Simon Institut d'Aeronomie Spatiale de Belgique Belgium

Susan Solomon NOAA Aeronomy Laboratory US

Petteri Taalas Finnish Meteorological Institute Finland

David Thompson Colorado State University US

Jan C. van der Leun Ecofys The Netherlands

Guus J.M. Velders National Institute of Public Health and the Environment (RIVM) The Netherlands

Jean Verdebout European Commission Joint Research Centre Italy

Robert T. Watson World Bank US

Elizabeth C. Weatherhead NOAA Air Resources Laboratory US

Ray F. Weiss University of California at San Diego/Scripps Institution US

of Oceanography

Vladimir Yushkov Central Aerological Observatory Russia

Rodolphe J. Zander Université de Liège Belgium

Tong Zhu Peking University China

AUTHORS, CONTRIBUTORS, AND REVIEWERS
A.13

Sponsoring Organizations and Liaisons

Michael H. Proffitt World Meteorological Organization Switzerland

Marco González United Nations Environment Programme Kenya

Daniel L. Albritton National Oceanic and Atmospheric Administration US

Michael J. Kurylo National Aeronautics and Space Administration US

Georgios T. Amanatidis European Commission Belgium

Coordinating Editor

Christine A. Ennis NOAA Aeronomy Laboratory/CIRES US

Technical Editing

Christine C. Sweet Science and Technology Corporation US

Christine A. Ennis NOAA Aeronomy Laboratory/CIRES US

Publication Design and Layout

Debra Dailey-Fisher (Lead) NOAA Aeronomy Laboratory US

Albert D. Romero (Consulting and Support) NOAA Mountain Administrative Support Center US

Dennis Dickerson (Graphics Design, “Twenty Questions”) Concepts 3 US

Editorial Assistance

Jeanne S. Waters NOAA Aeronomy Laboratory US

Barbara A. Keppler NOAA Aeronomy Laboratory US

Conference Coordination and Documentation

Michael H. Proffitt World Meteorological Organization Switzerland

Marie-Christine Charriére World Meteorological Organization Switzerland

Christine A. Ennis NOAA Aeronomy Laboratory/CIRES US

Jeanne S. Waters NOAA Aeronomy Laboratory US

Kathy A. Thompson Computer Sciences Corporation US

Conference Support

Jeanne S. Waters NOAA Aeronomy Laboratory US

Kathy A. Thompson Computer Sciences Corporation US

Catherine A. Burgdorf NOAA Aeronomy Laboratory/CIRES US

Debra Dailey-Fisher NOAA Aeronomy Laboratory US

Rose M. Kendall Computer Sciences Corporation US

Randy Soderholm Computer Sciences Corporation US

Chantal Renaudot World Meteorological Organization US

Barbara A. Keppler NOAA Aeronomy Laboratory US

Suzie Milano-Schoser NOAA Aeronomy Laboratory US

Gordon McFadyen Scottish Environment Protection Agency UK

Neil R.P. Harris European Ozone Research Coordinating Unit UK

Computing and Networking Support

Catherine A. Burgdorf NOAA Aeronomy Laboratory/CIRES US

Walter J. Harrop NOAA Aeronomy Laboratory US

Gabrielle Accatino NOAA Aeronomy Laboratory US

Ken Jamieson NOAA Aeronomy Laboratory US

Richard J. Tisinai NOAA Aeronomy Laboratory US
B.1

APPENDIX B

MAJOR ACRONYMS AND ABBREVIATIONS

AASE Airborne Arctic Stratospheric Expedition

ADEOS Advanced Earth Observing Satellite

AER Atmospheric and Environmental Research, Inc. (United States)

AFEAS Alternative Fluorocarbons Environmental Acceptability Study

AGAGE Advanced Global Atmospheric Gases Experiment

AI aerosol index

ALE Atmospheric Lifetime Experiment

AO Arctic Oscillation

AOD aerosol optical depth

ATMOS Atmospheric Trace Molecule Spectroscopy

AVHRR Advanced Very High Resolution Radiometer

AWI Alfred Wegener Institute (Germany)

BDC Brewer-Dobson circulation

BL boundary layer

BUV Backscatter Ultraviolet (spectrometer)

CARIBIC Civil Aircraft for Regular Investigation of the Atmosphere Based on an Instrument Container

CARMA Community Aerosol and Radiation Model for Atmospheres

CCSR Center for Climate System Research (University of Tokyo)

CFC chlorofluorocarbon

CICERO Centre for International Climate and Environmental Research, Universitetet I Oslo (Norway)

CIE Commission Internationale de l’Éclairage (France)

CIRES Cooperative Institute for Research in Environmental Sciences (United States)

CLAES Cryogenic Limb Array Etalon Spectrometer

CLAMS Chemical Lagrangian Model of the Stratosphere

CMAM Canadian Middle Atmosphere Model

CMDL Climate Monitoring and Diagnostics Laboratory (NOAA)

CNR-IFA Consiglio Nazionale della Richerche-Instituto di Fisica dell’Atmosfera (Italy)

CNRM Centre National de Recherches Météorologiques (France)

CNRS Centre National de la Recherche Scientifique (France)

CONICET Consejo de Investigaciones Científicas y Técnicas

CPC Climate Prediction Center (NCEP)

CSIRO Commonwealth Scientific and Industrial Research Organisation (Australia)

CTM chemical transport model

DLR Deutschen Zentrum für Luft- und Raumfahrt (Germany)

DMS dimethyl sulfide

DMSO dimethyl sulfoxide

DOAS Differential Optical Absorption Spectroscopy

DSCD differential slant column density

DSIR Department of Scientific and Industrial Research (South Africa)

DU Dobson unit

EC European Commission

ECC electrochemical concentration cell (ozonesonde)
ACRONYMS

B.2

ECD electron capture detector

ECHAM European Centre Hamburg Model

ECMWF European Centre for Medium-Range Weather Forecasts (United Kingdom)

EECl effective equivalent chlorine

EESC effective equivalent stratospheric chlorine

ENSO El Niño-Southern Oscillation

EP Earth Probe

EPA Environmental Protection Agency (United States)

ER-2 Earth Resources-2 (aircraft)

ERBE Earth Radiation Budget Experiment

ERS-2 European Remote Sensing-2 (satellite)

ETH Eidgenössische Technische Hochschule (Swiss Federal Institute of Technology) (Switzerland)

EU European Union

EUVDB European Ultraviolet Database

FC fluorocarbon

FDH fixed dynamical heating (model)

FSSP forward scattering spectrometer probe

FTIR Fourier transform infrared

FUB Freie Universität Berlin (Germany)

GAGE Global Atmospheric Gases Experiment

GAW Global Atmosphere Watch

GCM general circulation model

GC-MS gas chromatograph-mass spectrometer

GFDL Geophysical Fluid Dynamics Laboratory (NOAA)

GHG greenhouse gas

GISS Goddard Institute for Space Studies (NASA)

GOME Global Ozone Monitoring Experiment

GRIPS GCM-Reality Intercomparison Project for SPARC

GSFC Goddard Space Flight Center (NASA)

GWD gravity wave drag

GWP Global Warming Potential

HALOE Halogen Occultation Experiment

HBFC hydrobromofluorocarbon

HC hydrocarbon

HCFC hydrochlorofluorocarbon

HFC hydrofluorocarbon

HFE hydrofluorinated ether or hydrofluoroether

hPa hectoPascal

IASB Institut d’Aéronomie Spatiale de Belgique (Belgium)

IFU Institute for Atmospheric Environmental Research (Germany)

IGAC International Global Atmospheric Chemistry

ILAS Improved Limb Atmospheric Spectrometer

IPCC Intergovernmental Panel on Climate Change

IR infrared

IROE-CNR Instituto di Riccrea sulle Onde Elettromagneticher del Consiglio Nazionale della Richerche (Italy)

ISAMS Improved Stratospheric and Mesospheric Sounder

ISCCP International Satellite Cloud Climatology Project

IUPAC International Union of Pure and Applied Chemistry
ACRONYMS

JMA Japan Meteorological Agency (Japan)

JPL Jet Propulsion Laboratory (NASA)

JRC Joint Research Centre (Italy)

KASIMA Karlsruhe Simulation Model of the Middle Atmosphere

KNMI Koninklijk Nederlands Meteorologisch Instituut (The Netherlands)

LaRC Langley Research Center (NASA)

LCTM Lagrangian chemical transport model

LLNL Lawrence Livermore National Laboratory (United States)

LMS lowermost stratosphere

LS lower stratosphere

M3 Meteor-3 (satellite)

MAECHAM/CHEM Middle Atmosphere European Centre Hamburg Model with Chemistry

MASP Multiangle Aerosol Spectrometer Probe

MATCH Model of Atmospheric Transport and Chemistry

MBL marine boundary layer

MD mass deficiency

MIT Massachusetts Institute of Technology (United States)

MLS Microwave Limb Sounder

MODIS Moderate Resolution Imaging Spectroradiometer

MOZAIC Measurement of Ozone and Water Vapor by Airbus In-Service Aircraft

MOZART2 Model for Ozone and Related Chemical Tracers, version 2

MPAE Max-Planck-Institut für Aeronomie (Germany)

MPIC Max-Planck-Institut für Chemie (Germany)

MSA methanesulfonic acid

MSC Meteorological Service of Canada

MSU Microwave Sounding Unit

N7 Nimbus-7 (satellite)

NAD nitric acid dihydrate

NAM NH annular mode

NAO North Atlantic Oscillation

NASA National Aeronautics and Space Administration (United States)

NAT nitric acid trihydrate

NCAR National Center for Atmospheric Research (United States)

NCEP National Centers for Environmental Prediction (NOAA)

NDSC Network for the Detection of Stratospheric Change

NH Northern Hemisphere

NIES National Institute for Environmental Studies (Japan)

NIST National Institute of Standards and Technology (formerly NBS, United States)

NIWA National Institute of Water and Atmospheric Research (New Zealand)

NMHC nonmethane hydrocarbon

NOAA National Oceanic and Atmospheric Administration (United States)

n-PB n-propyl bromide

NPL National Physical Laboratory (United Kingdom)

NPLS nonparametric least-square fits

NSF National Science Foundation (United States)

NWP numerical weather prediction

NWS National Weather Service (NOAA)

B.3
ACRONYMS

ODP Ozone Depletion Potential

ODS ozone-depleting substance

OGI Oregon Graduate Institute (United States)

OHP Observatoire de Haute-Provence (France)

OMS Observations of the Middle Stratosphere

OPC optical particle counter

PAN peroxyacetyl nitrate

PAUR Photochemical Activity and Solar Ultraviolet Radiation

PEM-Tropics Pacific Exploratory Mission in the Tropical Pacific

PFC perfluorocarbon

PGI product gas injection

POAM Polar Ozone and Aerosol Measurement

POLARIS Photochemistry of Ozone Loss in the Arctic Region in Summer

ppb parts per billion

ppbv parts per billion by volume

ppm parts per million

ppmv parts per million by volume

ppt parts per trillion

pptv parts per trillion by volume

PSC polar stratospheric cloud

PV potential vorticity

PWD planetary-wave drag

QBO quasi-biennial oscillation

RAF radiation amplification factor

RDF reverse domain filling

RIVM Rijksinstituut voor Volksgezondheid en Milieu (National Institute of Public Health and the

Environment) (The Netherlands)

rms root mean square

RT radiative transfer

SAD surface area density

SAGE Stratospheric Aerosol and Gas Experiment

SAM SH annular mode

SAM Stratospheric Aerosol Measurement

SAOZ Système d’Analyse par Observation Zénithale

SBDART Santa Barbara Discrete-ordinate Hemispheric Radiative Transfer

SBUV/SBUV2 Solar Backscatter Ultraviolet (spectrometer)

SCD slant column density

SeaWiFS Sea-viewing Wide Field-of-view Sensor

SEFDH seasonally evolving fixed dynamical heating (model)

SGI source gas injection

SH Southern Hemisphere

SIO Scripps Institution of Oceanography (United States)

SOLVE SAGE III Ozone Loss and Validation Experiment

SPARC Stratospheric Processes and Their Role in Climate (WCRP)

SPOT Satellite Pour l’Observation de la Terre (satellite)

SRES Special Report on Emissions Scenarios (IPCC)

SRM solid rocket motor

SSA stratospheric sulfate aerosol

SSU Stratospheric Sounding Unit

B.4
ACRONYMS

STAR System for Transfer of Atmospheric Radiation

STE stratosphere-troposphere exchange

STRAT Stratospheric Tracers of Atmospheric Transport

STREAM Stratosphere-Troposphere Experiments by Aircraft Measurements

STS supercooled ternary solution

STT secondary tropical tropopause

SUNY State University of New York (United States)

SUSPEN Standardization of Ultraviolet Spectroradiometry in Preparation of a European Network

SUVDAMA Scientific UV Data Management

SZA solar zenith angle

2-D two-dimensional

3-D three-dimensional

TAR Third Assessment Report (IPCC)

TEAP Technology and Economic Assessment Panel (UNEP)

THESEO Third European Stratospheric Experiment on Ozone

TIROS Television Infrared Observation Satellite

TOMS Total Ozone Mapping Spectrometer

TOVS TIROS Operational Vertical Sounder

TRACE-P Transport and Chemical Evolution over the Pacific 

TT tropical tropopause 

TTL tropical tropopause layer

TTT thermal tropical tropopause

TUV Tropospheric Ultraviolet Visible

UARS Upper Atmosphere Research Satellite

UCI University of California at Irvine (United States)

UEA University of East Anglia (United Kingdom)

UH University of Heidelberg (Germany)

UIUC University of Illinois at Urbana-Champaign (United States)

UK United Kingdom

UKMO United Kingdom Meteorological Office

ULAU Università degli Studi dell’Aquila (Italy)

UM unified model

UMETRAC Unified Model with Eulerian Transport and Chemistry

UN United Nations

UNEP United Nations Environment Programme

US upper stratosphere

USDA United States Department of Agriculture

UT University of Tokyo (Japan)

UT upper troposphere

UV ultraviolet

VOC volatile organic compound

WCRP World Climate Research Programme

WMGHG well-mixed greenhouse gas

WMO World Meteorological Organization

WOUDC World Ozone and Ultraviolet Radiation Data Centre (Canada)

B.5
C.1

APPENDIX C

MAJOR CHEMICAL FORMULAE AND NOMENCLATURE

FROM THIS ASSESSMENT

HALOGEN-CONTAINING SPECIES

Cl atomic chlorine Br atomic bromine

Cl2 molecular chlorine Br2 molecular bromine

ClO chlorine monoxide BrO bromine monoxide

ClOx chlorine radicals BrOx bromine radicals

OClO chlorine dioxide OBrO bromine dioxide

ClOO chloroperoxy radical

Cl2O2, ClOOCl dichlorine peroxide (ClO dimer)

ClONO chlorine nitrite, nitryl chloride

ClONO2, ClNO3 chlorine nitrate BrONO2, BrNO3 bromine nitrate

HCl hydrogen chloride (hydrochloric acid) HBr hydrogen bromide

HOCl hypochlorous acid HOBr hypobromous acid

Cly inorganic chlorine Bry inorganic bromine

CCly organic chlorine CBry organic bromine

BrCl bromine chloride BrOOCl bromochloroperoxide

F atomic fluorine I atomic iodine

F2 molecular fluorine I2 molecular iodine

IO iodine monoxide

OIO iodine dioxide

IONO2 iodine nitrate

HF hydrogen fluoride (hydrofluoric acid) HI hydrogen iodide

HOI hypoiodous acid

SF6 sulfur hexafluoride

SF5CF3 trifluoromethylsulfurpentafluoride

NF3 nitrogen trifluoride

HALOCARBONS

CHLOROFLUOROCARBONS (CFCS) HYDROCHLOROFLUOROCARBONS (HCFCS)

CFC-11 CCl3F HCFC-21 CHCl2F

CFC-12 CCl2F2 HCFC-22 CHClF2

CFC-13 CClF3 HCFC-31 CH2ClF

CFC-113 CCl2FCClF2 HCFC-123 CHCl2CF3

CFC-113a CCl3CF3 HCFC-124 CHClFCF3

CFC-114 CClF2CClF2 HCFC-141b CH3CCl2F

CFC-114a CCl2FCF3 HCFC-142b CH3CClF2

CFC-115 CClF2CF3 HCFC-225ca CHCl2CF2CF3

HCFC-225cb CHClFCF2CClF2

HCFC-243cc CH3CF2CFCl2
CHEMICAL FORMULAE

C.2

HYDROFLUOROCARBONS (HFCS)

HFC-23 CHF3 HFC-245cb CH3CF2CF3

HFC-32 CH2F2 HFC-245ca CH2FCF2CHF2

HFC-41 CH3F HFC-245ea CHF2CHFCHF2

HFC-125 CHF2CF3 HFC-245eb CH2FCHFCF3

HFC-134 CHF2CHF2 HFC-245fa CHF2CH2CF3

HFC-134a CH2FCF3 HFC-263fb CH3CH2CF3

HFC-143 CH2FCHF2 HFC-272ca CH3CF2CH3

HFC-143a CH3CF3 HFC-281ea CH3CHFCH3

HFC-152 CH2FCH2F HFC-365mfc CH3CF2CH2CF3

HFC-152a CH3CHF2 HFC-356mcf CH2FCH2CF2CF3

HFC-161 CH3CH2F HFC-356mff CF3CH2CH2CF3

HFC-227ea CF3CHFCF3 HFC-338pcc CHF2CF2CF2CHF2

HFC-236cb CH2FCF2CF3 HFC-43-10mee CF3CHFCHFCF2CF3

HFC-236ea CHF2CHFCF3 HFC-458mfcf CF3CH2CF2CH2CF3

HFC-236fa CF3CH2CF3 HFC-55-10mcff CF3CF2CH2CH2CF2CF3

HALONS

Halon-1202 CBr2F2 Halon-1301 CBrF3

Halon-1211 CBrClF2 Halon-2402 CBrF2CBrF2

CHLOROCARBONS BROMOCARBONS

CH3Cl methyl chloride, chloromethane CH3Br methyl bromide, bromomethane

CH2Cl2 methylene chloride, dichloromethane CH2Br2 methylene bromide, dibromomethane

CHCl3 chloroform, trichloromethane CHBr3 bromoform, tribromomethane

CCl4 carbon tetrachloride CH2BrCH2Br 1,2 dibromoethane

C2HCl3 trichloroethene, trichloroethylene CH3CH2CH2Br 1-bromopropane, n-propyl bromide, n-PB

C2Cl4 tetrachloroethene, perchloroethene (n-C3H7Br)

CH3CH2Cl chloroethane CH3C(O)CH2Br bromoacetone

CH2ClCH2Cl 1, 2 dichloroethane

CH3CCl3 methyl chloroform FLUOROCARBONS

C2H2Cl4 tetrachloroethane CF4 tetrafluoromethane, carbon tetrafluoride

C4Cl6 hexachlorobutadiene C2F6, CF3CF3 perfluoroethane

COCl2 phosgene, carbonyl chloride C3F8, CF3CF2CF3 perfluoropropane

c-C3F6 perfluorocyclopropane

IODOCARBONS C4F10 perfluorobutane

CH3I iodomethane, methyl iodide c-C4F8 perfluorocyclobutane

CH2I2 diiodomethane C5F12 perfluoropentane

CH3CH2I iodoethane, ethyl iodide C6F14 perfluorohexane

CH3CHICH3 2-iodopropane, isopropyl iodide COF2 carbonyl fluoride

(i-C3H7I) TFA trifluoroacetic acid (CF3COOH)

CH3CH2CH2I 1-iodopropane, n-propyl iodide CF3COF trifluoroacetyl fluoride

(n-C3H7I)

OTHERS

CHBr2Cl dibromochloromethane CF3I trifluoromethyl iodide, trifluoroiodomethane

CH2BrCl bromochloromethane CF3CF2I iodopentafluoroethane

CHBrCl2 bromodichloromethane COClF chlorofluorocarbonyl

CH2BrI bromoiodomethane CF3COCl trifluoroacetyl chloride

CHBrF2 bromodifluoromethane SF6 sulfur hexafluoride

CH2ClI chloroiodomethane SF5CF3 trifluoromethylsulfurpentafluoride
C.3

OTHER SPECIES

O atomic oxygen H atomic hydrogen

O(3P) atomic oxygen (ground state) H2 molecular hydrogen

O(1D) atomic oxygen (first excited state) OH hydroxyl radical

O2 molecular oxygen HO2 hydroperoxyl radical

O3 ozone H2O water

Ox odd oxygen (O, O(1D), O3) or H2O2 hydrogen peroxide

oxidant (O3 + NO2) HOx odd hydrogen (H, OH, HO2, H2O2)

HDO deuterated water

N atomic nitrogen HNO2, HONO nitrous acid

N2 molecular nitrogen HOONO pernitrous acid

N2O nitrous oxide HNO3 nitric acid

NO nitric oxide HNO4, HO2NO2 peroxynitric acid, pernitric acid

NO2 nitrogen dioxide CH3OONO2 methylperoxynitrate

NO3 nitrogen trioxide, nitrate radical PAN peroxyacetylnitrate (CH3C(O)OONO2)

N2O5 dinitrogen pentoxide RONO2 alkyl nitrates

NOx nitrogen oxides (NO + NO2) NAD nitric acid dihydrate (HNO3◊2H2O)

NOy total reactive nitrogen NAT nitric acid trihydrate (HNO3◊3H2O)

(usually includes NO, NO2, NO3, NAP nitric acid pentahydrate (HNO3◊5H2O)

N2O5, ClONO2, HNO4, HNO3)

S atomic sulfur H2S hydrogen sulfide

SO2 sulfur dioxide CS2 carbon disulfide

SO4 sulfate COS, OCS carbonyl sulfide

H2SO4 sulfuric acid

CH3SCH3 DMS, dimethyl sulfide

CH3S(O)CH3 DMSO, dimethyl sulfoxide

CH3SO3H MSA, methanesulfonic acid

C carbon atom

CO carbon monoxide

CO2 carbon dioxide

NMHC nonmethane hydrocarbon CH2O, HCHO formaldehyde

CH4 methane CH3COCH3 acetone

C2H6 ethane CH3OOH methyl hydroperoxide

C3H8 propane CH3COO methyl peroxy radical

C2H4 ethylene, ethene CH3C(O)OO acetyl peroxy radical

C2H2 acetylene, ethyne RO alkoxy radicals

RO2 organic peroxy radicals

CHEMICAL FORMULAE